Biogeochemistry of Marine Systems (Sheffield Biological Sciences S.)

Biogeochemistry of Marine Systems

Biological Sciences Series A series which provides an accessible source of information at research and professional level in chosen sectors of the biological sciences. Series Editors: Professor Jeremy A. Roberts, Plant Science Division, School of Biosciences, University of Nottingham. Professor Peter N.R. Usherwood, Molecular Toxicology Division, School of Biosciences, University of Nottingham. Titles in the series: Stress Physiology in Animals Edited by P.H.M. Balm Seed Technology and its Biological Basis Edited by M. Black and J.D. Bewley Leaf Development and Canopy Growth Edited by B. Marshall and J.A. Roberts Environmental Impacts of Aquaculture Edited by K.D. Black Herbicides and their Mechanisms of Action Edited by A.H. Cobb and R.C. Kirkwood The Plant Cell Cycle and its Interfaces Edited by D. Francis Meristematic Tissues in Plant Growth and Development Edited by M.T. McManus and B.E. Veit Fruit Quality and its Biological Basis Edited by M. Knee Pectins and their Manipulation Edited by G.B. Seymour and J.P. Knox Biogeochemistry of Marine Systems Edited by K.D. Black and G.B. Shimmield

Biogeochemistry of Marine Systems Edited by KENNETH D. BLACK Scottish Association for Marine Science Dunstaffnage Marine Laboratory Oban, UK and GRAHAM B. SHIMMIELD Director, Scottish Association for Marine Science Dunstaffnage Marine Laboratory Oban, UK

Blackwell Publishing

© 2003 by Blackwell Publishing Ltd Editorial Offices: Blackwell Publishing Ltd, 9600 Garsington Road, Oxford OX4 2DQ, UK Tel: +44 (0)1865 776868 Blackwell Publishing Asia Pty Ltd, 550 Swanston Street, Carlton, Victoria 3053, Australia Tel: +61 (0)3 8359 1011 ISBN 1–84127–327–9 Originated as Sheffield Academic Press Published in the USA and Canada (only) by CRC Press LLC, 2000 Corporate Blvd., N.W., Boca Raton, FL 33431, USA Orders from the USA and Canada (only) to CRC Press LLC USA and Canada only: ISBN 0–8493–2818–7 The right of the Author to be identified as the Author of this Work has been asserted in accordance with the Copyright, Designs and Patents Act 1988. All rights reserved. No part of this publication may be reproduced, stored in a retrieval system, or transmitted, in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, except as permitted by the UK Copyright, Designs and Patents Act 1988, without the prior permission of the publisher. This book contains information obtained from authentic and highly regarded sources. Reprinted material is quoted with permission, and sources are indicated. Reasonable efforts have been made to publish reliable data and information, but the author and the publisher cannot assume responsibility for the validity of all materials or for the consequences of their use.

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Contents

Preface Contributors 1 Mangroves of Southeast Asia

xii xv 1

M. HOLMER 1.1 1.2 1.3 1.4

Introduction Mangrove forest structure and function Water column biogeochemistry Organic matter sources in mangrove forests 1.4.1 Decomposition of detritus 1.5 Sediment biogeochemistry 1.5.1 Total microbial activity in mangrove sediments 1.5.2 Mineralization pathways in mangrove sediments 1.5.3 Phosphorus cycling 1.6 Factors influencing the biogeochemistry 1.6.1 Effect of forest type and age 1.6.2 Influence of macrofauna 1.6.3 Effect of seasonal variations on mangrove forest biogeochemistry 1.7 Sediment biogeochemistry and implications for mangrove vegetation 1.8 Biogeochemistry in mangroves affected by anthropogenic activities References 2 Coral reefs

1 2 3 7 9 11 12 15 23 25 25 26 28 29 31 34 40

M.J. ATKINSON and J.L. FALTER 2.1 Introduction 2.2 Coral reef morphology and zonation 2.3 Basic biogeochemistry 2.3.1 Carbon 2.3.2 Dissolved organic matter 2.3.3 Nitrogen 2.3.4 Phosphorus

40 41 43 43 47 48 50

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CONTENTS

2.3.5 Silica 2.3.6 Iodine 2.4 Interstitial geochemistry and hydrology of coral reef frameworks 2.5 Mass transfer-limited biogeochemical rates 2.6 Coral growth in high nutrient water 2.7 Measurement techniques 2.8 Summary statements References 3 Fjords

51 51 52 54 56 58 59 59 65

J.M. SKEI, B. MCKEE and B. SUNDBY 3.1 Introduction 3.1.1 Definition and origin of fjords 3.1.2 The public and scientific interest in fjords 3.2 Sediment diagenesis in oxic fjords 3.2.1 The Saguenay Fjord 3.2.2 Sedimentation 3.2.3 Composition of the rapidly deposited layers 3.2.4 Sulfate reduction and sulfide accumulation 3.2.5 Mercury diagenesis 3.2.6 Phosphorus and arsenic geochemistry 3.2.7 Non-steady-state diagenesis 3.3 Elemental cycling in anoxic waters 3.3.1 Chemical tracers 3.3.2 Cycling of carbon and nutrients 3.3.3 Trace element and radionuclide cycling 3.3.4 Fe–S systematics 3.3.5 Sulfate reduction and methane oxidation 3.3.6 Elemental cycling in sediments underlying anoxic waters 3.3.7 Preservation of organic matter References 4 The Eastern Mediterranean

65 66 67 69 69 69 70 70 73 73 75 76 76 77 78 81 83 84 85 86 91

MICHAEL D. KROM, STEVEN GROOM and TAMAR ZOHARY 4.1 Introduction 4.2 History of the Mediterranean basin 4.3 Basic description of the bathymetry and physical oceanography of the Eastern Mediterranean 4.3.1 Bathymetry

91 92 93 93

CONTENTS

4.3.2 Physical circulation of the Eastern Mediterranean 4.3.2.1 Formation of LDW 4.3.2.2 Formation of LIW 4.3.3 Recent water mass changes in the Eastern Mediterranean 4.3.4 Current patterns 4.4 Nutrients and chlorophyll distribution across the Eastern Mediterranean 4.4.1 General comments 4.4.2 Seasonal distributions 4.4.2.1 Winter 4.4.2.2 Spring into summer 4.4.3 Nutrient distribution below the nutricline 4.5 Total chlorophyll distribution and characteristics 4.5.1 Light penetration 4.5.2 Species composition 4.5.2.1 The prochlorophytes 4.5.2.2 The unicellular cyanobacteria 4.5.2.3 The eukaryotes 4.5.2.4 Heterotrophic bacteria 4.6 Primary production 4.6.1 Gradient in biomass and productivity from coastal waters to the open sea 4.7 Effects of mesoscale features on nutrient and chlorophyll distribution and phytoplankton productivity 4.7.1 Biogeochemical processes in mesoscale features 4.7.1.1 Rhodes cold-core (cyclonic) eddy 4.7.1.2 Cyprus warm-core (anti-cyclonic) eddy 4.7.1.3 Effects of other mesoscale features 4.8 Seasonal changes in phytoplankton biomass as detected by remote sensing 4.9 Nutrient limitation in the Eastern Mediterranean 4.10 Magnitude of man-induced changes in nutrient inputs and their possible effects on the Eastern Mediterranean 4.11 Summary and conclusions Acknowledgements Glossary References 5 The Arctic seas

vii 93 94 95 96 97 98 98 98 98 100 100 100 102 102 102 103 104 104 105 107 108 108 108 112 113 113 116 118 120 121 122 122 127

MICHAEL L. CARROLL and JOLYNN CARROLL 5.1 5.2

Summary Main features

127 128

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CONTENTS

5.2.1 Water masses 5.2.2 Continental shelves 5.2.3 Sea ice 5.3 Biogeochemical cycles and ecological processes 5.4 Environmental changes 5.4.1 Climate variability 5.4.2 Long-term climate change 5.4.3 Ozone and ultraviolet radiation 5.4.4 Contaminants 5.5 Natural resources and ecological services 5.5.1 Indigenous people 5.5.2 Non-indigenous regional populations 5.5.3 National/international/global users Acknowledgements References 6 The Arabian Sea

128 131 132 133 138 138 140 143 143 144 144 145 146 147 147 157

S.W.A. NAQVI, HEMA NAIK and P.V. NARVEKAR 6.1 6.2 6.3 6.4

Introduction Geographical setting Climate and circulation Nutrients and primary production 6.4.1 Subsurface nutrient trap 6.4.2 Primary productivity 6.4.3 New production 6.4.4 Phytoplankton composition and size distribution 6.4.5 Chlorophyll and POC 6.4.6 Effect of changes in mixed layer depth 6.5 Heterotrophic biomass and production 6.5.1 Heterotrophic bacteria 6.5.2 Nano- and microheterotrophs 6.5.3 Mesozooplankton 6.6 Food web structure and export of material to the deep sea 6.6.1 Phytoplankton growth and mortality 6.6.2 Particle fluxes to deep sea 6.6.3 Role of Arabian Sea as a source or sink of carbon dioxide (CO2) 6.7 Oxygen-deficient zones 6.7.1 Denitrification 6.7.2 Intermediate nepheloid layer 6.7.3 Other redox-sensitive elements 6.7.4 Biological effects

157 157 158 163 163 164 166 167 171 172 174 174 175 177 179 179 180 185 185 186 191 192 193

CONTENTS

6.8 Benthic processes References 7 The northeastern Pacific abyssal plain

ix 195 198 208

ANGELOS K. HANNIDES and CRAIG R. SMITH 7.1 Introduction 7.2 Key habitat parameters of deep seafloor communities 7.2.1 Key habitat parameters 7.2.1.1 Substratum type 7.2.1.2 Near-bottom currents 7.2.1.3 Bottom-water oxygen 7.2.1.4 Sinking POC flux 7.2.1.5 Redox conditions 7.2.2 Variation of key habitat parameters in the northeastern Pacific abyssal plain 7.2.2.1 Sediment types 7.2.2.2 Near-bottom currents and oxygen concentrations 7.2.2.3 POC flux and redox conditions 7.3 Northeastern Pacific abyssal zones 7.3.1 The eutrophic equatorial abyss 7.3.2 The mesotrophic (sub-equatorial) abyss 7.3.3 The oligotrophic central gyre abyss 7.4 Sensitivity and resilience to natural and anthropogenic change 7.4.1 General thoughts 7.4.2 Potential sensitivity and resilience to specific changes 7.4.2.1 Climate variation in the equatorial and North Pacific 7.4.2.2 Global increase in atmospheric greenhouse gases and temperatures 7.4.2.3 Manganese nodule mining 7.4.2.4 Iron fertilization 7.5 Concluding remarks Acknowledgments References 8 Deep-sea hydrothermal vents and cold seeps

208 208 209 209 210 211 211 211 212 212 212 213 213 214 217 218 220 220 221 221 223 226 229 230 231 231 238

RICHARD J. LÉVEILLÉ and S. KIM JUNIPER 8.1 Introduction 8.1.1 Deep-sea hydrothermal vents and cold seeps 8.1.2 Life at hydrothermal vents and cold seeps 8.1.3 Scope of this chapter

238 238 238 240

x

CONTENTS

8.2 Deep-sea hydrothermal vents 8.2.1 Distribution and general characteristics 8.2.1.1 Geochemical fluxes of gases and elements from hydrothermal vents 8.2.1.2 Off-axis diffuse flow versus axial venting 8.2.2 Subsurface biosphere at mid-ocean ridges 8.2.2.1 Evidence for a subsurface biosphere at deep-sea hydrothermal vents 8.2.2.2 Biogeochemical interactions in subsurface environments 8.2.3 Seafloor microbe-mineral interactions at hydrothermal vents 8.2.3.1 Microbial distribution and activity at vents 8.2.3.2 Biomineralisation at vents 8.2.3.3 Fossilisation of microbes at vents 8.2.3.4 Bacterial weathering of sulphides 8.2.4 Biogeochemical interactions in hydrothermal plumes 8.2.4.1 General features of hydrothermal plumes 8.2.4.2 Microbial ecology of hydrothermal plumes 8.2.4.3 Microbial productivity and organic carbon in plumes 8.2.4.4 Biogeochemical interactions in plumes 8.2.5 Biogeochemistry of off-axis vents and seafloor basalt 8.2.5.1 Off-axis vents 8.2.5.2 Seafloor basalts 8.3 Deep-sea cold seeps 8.3.1 Distribution, occurrences and general characteristics 8.3.1.1 Gas hydrates 8.3.1.2 Geochemical fluxes 8.3.2 Biogeochemistry of seep sediment pore fluids 8.3.2.1 Methanogenesis 8.3.2.2 Anaerobic sulphate reduction 8.3.2.3 Aerobic microbial oxidation of sulphide and methane 8.3.2.4 Anaerobic oxidation of methane 8.3.3 Microbial carbonates 8.4 Stability and perturbations of seafloor hydrothermal vent and cold seep systems 8.4.1 Geological stability of vents and seeps 8.4.2 Future perturbations related to resource extraction 8.4.2.1 Hydrothermal sulphides 8.4.2.2 Cold seeps 8.4.3 Response of cold seeps and gas hydrates to global warming

241 241 243 245 246 246 251 252 252 256 260 260 261 261 262 263 264 265 265 266 267 267 268 269 270 271 271 272 273 274 276 276 277 277 278 278

CONTENTS

8.5 Future work 8.6 Conclusion References 9 Influence of nutrient biogeochemistry on the ecology of northwest European shelf seas

xi 279 282 282

293

PAUL TETT, DAVID HYDES and RICHARD SANDERS 9.1 Introduction 9.2 Nutrient cycles 9.2.1 Macronutrient element availability 9.2.2 Sources of macronutrients 9.2.3 Sinks of macronutrients 9.2.4 Observed distributions of macronutrient concentrations and ratios 9.2.5 Iron 9.3 Plankton biogeochemistry 9.3.1 Taxonomy and life forms in the plankton 9.3.2 Theories of floristic composition 9.3.2.1 Light-nutrient-mixing explanations 9.3.2.2 Biogeochemical controls 9.3.2.3 Ecological controls 9.3.3 Variation in nutrient element ratios and its explanation in terms of plankton biochemistry 9.3.4 Quantitative theory for nutrient element ratios 9.3.5 Differences in abilities to assimilate different nutrients 9.3.6 Theoretical conclusions 9.4 Effects of ambient nutrient ratios on plankton 9.4.1 Introduction 9.4.2 Time series: Helgoland and the German Bight 9.4.3 Mesocosm and other competition experiments 9.4.4 Observations at sea 9.5 Discussion and conclusions 9.5.1 Introduction 9.5.2 Do high ambient N:Si ratios favour flagellates? 9.5.3 Do non-Redfield ambient N:P ratios perturb pelagic ecosystems? 9.5.4 The possibility of iron limitation in shelf seas 9.5.5 Trophic consequences of ratio changes – a Panglossian conclusion? 9.5.6 A flexible Redfield ratio? Dedication References Index

293 294 294 296 298 300 301 303 303 308 309 310 312 315 320 325 327 327 327 329 332 337 341 341 342 344 345 345 347 350 351 364

Preface

Marine biogeochemistry is a broad, interdisciplinary subject overlapping a range of other disciplines such as marine chemistry, geochemistry, ecology, physiology and oceanography, but in its own right it has become pivotal to progress in marine research in recent years. As a key component of the ‘earth system’, marine biogeochemistry interfaces directly with terrestrial, atmospheric and geological sciences. A working definition of the subject might be ‘the processing, recycling, storage, transport and loss of chemical components within the marine environment, mediated by biological processes’. We are said to be leaving the Holocene and entering the ‘Anthropocene’ where mankind’s cumulative impacts have significant and measurable effects on the biosphere. Biogeochemistry lies at the heart of studies on the functioning of marine provinces or types – collectively here referred to as ‘systems’ – that are crucial to understanding and predicting global change and its consequences. In the context of this great environmental and societal impact, it is the varying consequences of the same biogeochemical processes operating in marine systems under different forcing parameters that make biogeochemistry such a diverse and fascinating field. Over the past two decades, much has been learned about the biogeochemical functioning of marine systems from large-scale, multi-partner, international and national research programmes such as are supported by the International Geosphere–Biosphere Program (IGBP), Scientific Committee on Ocean Research (SCOR), Joint Global Ocean Flux Study (JGOFS) and its regional studies, and Land–Ocean Interactions in the Coastal Zone (LOICZ). In the UK, the supporting national programmes were the Biogeochemical Ocean Flux Study (BOFS) and Land–Ocean Interaction Study (LOIS) programmes. These have been undertaken typically on ‘process’ research cruises, where the focus has been on quantifying fluxes of key components (particularly carbon) within the ocean as well as between the ocean and its boundaries (land, sediments and atmosphere). Whilst considerable information continues to be derived from such studies, the expense of such undertakings, together with the relatively low temporal and spatial coverage offered, has led some biogeochemists to develop and use new methods of data collection. These include satellite and airborne remote sensing, benthic landers, autonomous underwater vehicles, and moored and drifting sensor packages with intelligence. Many of these systems have been developed for open ocean deployment, but they are also becoming modified

PREFACE

xiii

for use in shallow, coastal locations. We can expect to see further developments, together with new and more robust sensors and increased data collection and transmission capacity, leading to great improvements in knowledge, operating in a synoptic fashion (for example, the new ARGO programme of drifting subsurface floats across the ocean basins). Modelling has become ubiquitous in biogeochemistry, as in marine science more generally. Significant computing power is now available for the nesting of biogeochemical models within physical oceanographic models with high spatial resolution. Not only does this allow the generalisation of measurements made at a point in space and time, but it allows assessments and comparisons of the relative sensitivities of systems to external changes such as are caused, for example, by increased temperature, deepwater trawling or hydrocarbon exploration. This volume provides an overview of recent research on the biogeochemistry of a diverse range of complex marine systems, each of great importance to the ‘earth system’ but for varying reasons. The systems were chosen to emphasise different forcing factors, thus offering interesting contrasts. We have been fortunate that the chapter authors reflect the diversity of academic backgrounds that typifies biogeochemical research and that they have approached their tasks from varied perspectives. Thus, the repetition of basic concepts between chapters is kept to a minimum. The book will be read by researchers and advanced students of biogeochemistry, who will enjoy the contrasts between the systems chosen, and by workers in related areas of earth science, who will find that it provides a useful point of access to the primary literature across a broad range of marine biogeochemical processes. The first chapter deals with mangroves – key providers of biogeochemical services in large areas of tropical coastal areas that are under threat from insensitive development pressures. We stay in the tropics to consider coral biogeochemistry – also under threat in many areas from a combination of climate change, eutrophication, tourism and destructive fishing – before moving to fjords – the main interface between land and ocean in high latitudes. The eastern Mediterranean continues to attract considerable attention as a highly nutrient-poor and low productivity area, in stark contrast to the Arctic, which is light-limited during the winter months, highly productive in the summer, shows strong benthic–pelagic coupling over the shelf areas and has a productive community associated with the underside of sea ice. In the Arabian Sea, the biogeochemical system is under the control of large-scale, monsoon-linked circulation reversals with a pronounced oxygen minimum zone, and again this is an area under continuing scrutiny for its potential role in the nitrogen biogeochemical cycle. The sediments of the northeast Pacific abyss are dominated by a strong latitudinal gradient of carbon input across the equatorial divergence that has a profound effect on benthic productivity. Even in an area so remote from land, the threat of anthropogenic disturbance in the form of metal nodule mining is very real.

xiv

PREFACE

The penultimate chapter deals with the unique biogeochemistry of the hot and cold vents associated with plate tectonics – largely unknown until relatively recently and now thought to be of great significance in maintaining the basic chemistry of sea water. The book is completed by a modelling section, in which the ecology of planktonic organisms is examined in biogeochemical terms with an emphasis on modelling the interactions between pelagic chemistry and ecology in shelf seas, where significant recycling of sedimentary nutrients supplements direct terrestrial inputs. We are grateful to all the authors who have contributed to this volume. Each author has original insights and approaches and so each chapter is fresh and the whole volume novel and readable. We are particularly indebted to Graeme MacKintosh and David McDade at Blackwell Publishing, who have offered every support and encouragement to this project. Kenneth D. Black Graham B. Shimmield

Contributors

M.J. Atkinson

University of Hawaii SOEST, Hawaii Institute of Marine Biology, PO Box 1346, Kaneohe Bay Hawaii 96744 JoLynn Carroll Akvaplan-niva, Polar Environmental Centre, 14 Hjalmar Johansensgate, N-9296 Tromsø, Norway Michael L. Carroll Akvaplan-niva, Polar Environmental Centre, 14 Hjalmar Johansensgate, N-9296 Tromsø, Norway J.L. Falter University of Hawaii, SOEST, Hawaii Institute of Marine Biology, PO Box 1346, Kaneohe Bay Hawaii 96744 Steven Groom Plymouth Marine Laboratory, Prospect Place, Plymouth PL1 3DH Angelos K. Hannides Department of Oceanography, SOEST, University of Hawaii at Manoa, 1000 Pope Road, Marine Sciences Building, Honolulu, HI 96822 M. Holmer Institute of Biology, University of Southern Denmark, Campusvej 55, DK-5230 Odense, Denmark David Hydes Southampton Oceanography Centre, European Way, Southampton SO14 32H S. Kim Juniper Centre GEOTOP-UQÀM-McGill, Université du Québec à Montréal, C.P. 8888 Succ. Centre-Ville, Montréal, Québec, Canada, H3C 3P8 Michael D. Krom School of Earth Sciences, Leeds University, Leeds LS2 9JT Richard J. Léveillé Centre GEOTOP-UQÀM-McGill, Université du Québec à Montréal, C.P. 8888 Succ. Centre-Ville, Montréal, Québec, Canada, H3C 3P8 B. McKee Tulane University, New Orleans, USA Hema Naik National Institute of Oceanography, Dona Paula, Goa 403 004, India S.W.A. Naqvi National Institute of Oceanography, Dona Paula, Goa 403 004, India P.V. Narvekar National Institute of Oceanography, Dona Paula, Goa 403 004, India

xvi

Richard Sanders J.M. Skei Craig R. Smith

B. Sundby Paul Tett Tamar Zohary

CONTRIBUTORS

Southampton Oceanography Centre, European Way, Southampton SO14 32H Norwegian Institute for Water Research, Oslo, Norway Department of Oceanography, SOEST, University of Hawaii at Manoa, 1000 Pope Road, Marine Sciences Building, Honolulu, HI 96822 ISMER, Université du Québec à Rimouski and McGill University, Montreal, Canada School of Life Sciences, Napier University, 10 Colinton Road, Edinburgh, EH10 5DT Kinneret Limnological Laboratory, Israel Oceanographic and Limnological Research Ltd, PO Box 447, Migdal 14950, Israel

1

Mangroves of Southeast Asia M. Holmer

1.1

Introduction

The last decade has contributed significantly to the development of research on the biogeochemistry of tropical mangrove ecosystems. Also obvious during the last decade is the continued and dramatic destruction of natural tropical mangrove resources in Southeast Asia (Fig. 1.1). The naturally high productivity of tropical mangroves has traditionally been exploited for a wide variety of purposes, both as sources of forestry or fisheries products and they have also been used for human settlement (Hatcher et al., 1989; Platong, 1998). More recently, Southeast Asian mangroves are being extensively cleared for the construction of aquaculture ponds for prawn production (Primavera, 1993). This and other consumptive uses have been estimated to cause an annual reduction of 1% of the world’s tropical mangrove resources (Hatcher et al., 1989). Southeast Asian mangrove forests are declining at alarming rates, due to the increasing demand for land to be allocated to food, industrial production and urban settlements (Kautsky et al., 2000). More than half of the 367 900 ha of mangroves that was present in Thailand in 1961 had already been converted to prawn farms or for various other uses by 1989 (Aksornkoae, 1993), and the mangrove area was further reduced by 81 000 ha in 1996 (Platong, 1998). Mangrove forests in many other Southeast Asian countries also face the same rate of destruction, e.g. Vietnam (Kautsky et al., 2000). Changes in land use in Southeast Asia have resulted in high soil erosion rates and have yielded a major increase in transport of eroded sediments to the coastal zone. The long-term impact and the ramifications of anthropogenic disturbance, such as pollutant discharge, on the biogeochemistry of tropical mangrove forests and their associated near-shore habitats are poorly known. The aim of this chapter is to review recent findings on aspects of the biogeochemistry of Southeast Asian mangrove forests, focusing on physical and biotic processes determining the cycling of elements in mangroves. The chapter will discuss the significance of anthropogenic activities for the biogeochemical cycling of nutrients, which deserves particular attention in the study of Southeast Asian mangrove forests.

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Fig. 1.1 Areas of mangrove (solid line) and mangrove disturbance (hatched area) in the tropics.

1.2

Mangrove forest structure and function

Earlier models of tropical mangrove forests proposed overwhelming influences by physical forces and processes such as the tidal regime or geomorphology on ecosystem structure and function. Mangrove ecosystem development was depicted as successional systems, where the presence of plants themselves had a significant impact on the physical environment; and such impact culminated in an alternation of growth conditions favoured by different species in time. In addition, recent findings (especially from the Indo-Pacific mangrove forests) suggest considerable influence of biotic agents and processes such as sesarmine crab feeding and bioturbation activities in shaping the ecology of tropical mangrove forests (Twilley et al., 1997). Biotic influences on mangrove forest ecosystem structure and function are expected to be more important in systems with weak external forcing or high biodiversity. Climatic features, such as the timing of monsoon periods, can give rise to strong seasonality (Alongi & Sasekumar, 1992). Most productive mangrove ecosystems are highly effective sinks for nutrients essential to sustain high rates of plant growth, as evidenced from the fact that many such systems export refractory, particulate organic carbon but import most dissolved nutrient species (Alongi et al., 1992; Robertson et al., 1992; Hemminga et al., 1994; Rivera-Monroy et al., 1995; Alongi, 1996). Several studies have suggested that close couplings exist among benthic nutrient pools,

MANGROVES OF SOUTHEAST ASIA

3

microbes and mangrove trees, acting as mechanisms to maximize utilization and conservation of scarce nutrients (Boto & Wellington, 1983, 1984; Kristensen et al., 1995; Alongi, 1996).

1.3

Water column biogeochemistry

Mangrove forests are highly diversified due to the large variability, for example, in geomorphology and tidal activities, and a number of different functional types of mangroves have been described (Woodroffe, 1992). The hydrodynamics in some mangrove forests are strongly influenced by river inputs, whereas others are much more dominated by the ocean, as found for fringing mangrove forests. A large number of local factors may thus influence the water column processes, which makes it very difficult to provide a generalized description of the water biogeochemistry in mangrove forests. Mangrove creeks are considered as important routes for tidal exchange of dissolved and particulate matter between the forest environment and adjacent coastal waters (Wolanski et al., 1992; Hemminga et al., 1994; Rivera-Monroy et al., 1995). The residence time of the water in the creeks is usually a useful indicator of the biogeochemical fate of the compounds in the water column (Suraswadi et al., 2003), where long residence times allow for uptake of nutrients by the vegetation, e.g. phytoplankton and mangrove trees, while short residence times leads to a larger export of materials to the ocean (Thong et al., 1993). Hydrodynamics in the mangrove forest is controlled by tides, mangrove vegetation and geometry of the mangrove waterways (Hoguane et al., 1999). Friction from dense mangrove trees influences the tidal regime and causes tidal asymmetry (Suraswadi et al., 2003), and current velocities in a channel cross section have lateral or vertical variations due to channel geometry and bathymetry (Valle-Levinson & Atkinson, 1999). These variations in water velocity cause transverse and vertical shear stress, which are important for the mixing process in the creek water (Uncle et al., 1985). The residence time of water in mangrove forests is quite variable, determined by the forest topography, size and type, and thus hydrodynamics. It can vary from a few days in small fringe forests exposed to large tidal variation (Wattayakorn et al., 1990) to more than a month in large mangrove forests (Wolanski et al., 1990). Suraswadi et al. (2003) studied the hydrodynamics in a small mangrove forest in Thailand. The hydrodynamics was modified as a result of friction created by the mangrove vegetation, and these modifications resulted in strong ebb current, asymmetric flood and ebb tide and a time lag in the tidal phase between the upper and lower creek. The main creek was well mixed with a transient stratification during low tide and was completely mixed during high tide. This situation is similar to other mangrove estuaries (Wattayakorn et al., 1990; Wolanski et al., 1990). The hydrodynamics may be influenced by heavy rains; in this mangrove, rainfall caused an increased transport of salt water when the

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rainwater pushed the creek water outward, and the period and amount of rainfall must thus be considered when the flux of solutes from the mangrove forest is determined during the wet season. Nutrient levels in pristine tropical mangrove forests vary both in time and space as a result of differences in hydrodynamics, freshwater input, solar insolation, and productivity of phytoplankton and bacterioplankton (Ovalle et al., 1990; Alongi et al., 1992; Bava & Seralathan, 1998; Trott & Alongi, 1999; Ayakai et al., 2000). Mangrove creeks are, however, usually characterized by low nutrient concentrations due to a high capacity for retaining and recycling of nutrients within the system (Kristensen et al., 1995). Even in areas with high nutrient loading, e.g. due to urban settlement, nutrient levels are generally low (Harrison et al., 1997). Nutrient cycling in the water column is controlled by a large number of auto- and heterotrophic processes (Fig. 1.2), and the dissolved

Fig. 1.2 Nutrient transformation processes in the water column in mangrove forests.

MANGROVES OF SOUTHEAST ASIA

5

organic matter pool especially is considered to play an important role (Alongi et al., 1989; Bano et al., 1997; Ayukai et al., 1998). Leakage of nutrient rich water from the creek banks during low tides has been suggested as an important contributor to nutrients in the mangrove waterways, and was investigated in a mangrove forest in Thailand (Kristensen & Suraswadi, 2002). Water seeping from creek banks was only enriched in inorganic phosphates and was not considered as an important source of solutes to the waterways in this forest. The low nutrient concentrations affect the primary production in mangrove creeks, and the production has often been found to be strongly nutrient limited. Phosphorus was the limiting factor for phytoplankton growth in a mangrove creek in a tidal-dominated forest in Thailand, as indicated by a high molar ratio between dissolved nitrogen and phosphorus (~34–38) much higher than the Redfield ratio (Suraswadi et al., 2003). The degree of phosphorus limitation was less pronounced near small tributaries (~11–21), whereas the N:P ratio increased significantly (~192) near shrimp farm outlets – most likely in response to loss of dissolved nitrogen compounds from the farms (Burford & Longmore, 2001). The phytoplankton production is, however, not always nutrient limited. Due to the shallow water in many mangrove creeks and a rapid water flow in river-influenced mangrove forests the light penetration is often quite low and the phytoplankton production is just as often light limited (Harrison et al., 1997; Kristensen & Suraswadi, 2002). In a three year study of two tidal creeks on the Indus River in Pakistan, Harrison et al. (1997) found no limitation of the phytoplankton production by silicate or nitrogen, and phosphate was only limiting during large blooms. Conversely a large suspended load resulted in high light extinction coefficients and the 1% light depth ranged from <1 m to about 4 m. Consequently, primary productivity appeared to be light-limited throughout the year. There was no apparent seasonal cycle in primary productivity despite the variability in nutrient concentrations probably due to these poor light conditions. Gross primary production was also found to be only moderate in the water column in a Thailand mangrove forest dominated by tides with no pronounced spatial and temporal variations (Kristensen & Suraswadi, 2002). A large fraction of the inorganic carbon and nitrogen assimilated by the phytoplanktonic organisms was released in the form of DOC (50–90%) and DON (50–60%) indicating that primary producers were under stress or nutrient depleted. Phosphate and, occasionally, nitrogen appeared to be the limiting nutrients. However, rapid light attenuation in the turbid creek water rendered the entire water column strongly heterotrophic at all times. The high turbidity in mangrove waterways is caused by high loads of suspended matter. Suspended matter, which originates from the erosion of soil from forests and farmlands, is discharged in large quantities from rivers, especially in the humid tropics, where heavy rainfalls occur during the wet season. Resuspension in the shallow parts of the mangrove may also contribute significantly to suspended matter in the water column, whereas import from

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

the ocean plays a minor role. The suspended matter contains a variety of components including nutrients such as nitrogen and phosphorus. Most of the available phosphorus in mangrove waterways occurs in bound forms, rather than in free dissolved forms in the water. The role of suspended matter in nutrient enrichment in tropical mangroves is, however, poorly documented. One study of a mangrove forest in Malaysia showed that phosphorus and iron content in the suspended matter decreased linearly with increasing salinity, reflecting the process of phosphorus release into the mangrove waterways during transportation to the sea (Tanaka et al., 1998). Iron hydroxides display a strong affinity for phosphate and are considered to be primarily responsible for phosphate adsorption in oxidized environments (Sundby et al., 1992) such as the suspended matter in mangrove forests waterways. The average C:N:P ratio of the organic substances contained in the suspended matter in waterways in the Matang Mangrove Forest was, however, estimated to be 140:16:1, which is considerably different from the ratios for mangrove litter or terrestrial sources and is rather similar to the Redfield ratio, indicating a high contribution of living organisms to the suspended matter (Tanaka et al., 1998). It is thus very likely that the suspended matter undergo important transformations during the transportation in the mangrove forests. Except for a few recent studies (e.g. Bano et al., 1997; Harrison et al., 1997; Rivera-Monroy et al., 1998), there has been little focus on the role of microbial processes for transformations of dissolved and suspended compounds in the water column. These studies have indicated that much of the net nutrient generation within a mangrove forest originates from microbial mineralization of dissolved and particulate organic matter in tidal waters. Heterotrophic conditions in the mangrove waterways tend to result in accumulation of dissolved inorganic nutrients and dissolved organic matter, which eventually may be utilized by bacteria. Bacterial biomass and production was studied in three tidal creeks in the Indus River delta in Pakistan Bano et al. (1997), and they found a high bacterial production due to attached bacteria. The bacterial production was generally higher than the primary production, and the water columns were net heterotrophic. The microbial heterotrophs are considered to be largely supported by particulate and dissolved substrates derived from land run-off, tidal resuspension, mangrove root exudates and leachates from fallen leaves on the forest floor. In a mangrove forest in Thailand, it was, however, found that the pelagic heterotrophic community was supported primarily by exudates from primary producers (Kristensen & Suraswadi, 2002). In view of the low nutritional quality of the mangrove detritus, the production of phytoplankton and bacterial biomasses may represent important pathways for the synthesis of high quality biomass potentially available to the grazers in the mangrove creek systems (Bano et al., 1997). On the other hand, the heterotrophic conditions may also have negative consequences for the

MANGROVES OF SOUTHEAST ASIA

7

oxygen concentrations in the water column due to significant oxygen consumption particularly during night. Low oxygen concentrations have been measured in small tributaries in the high-intertidal forest, where the mixing of water is reduced due to low tidal influence (Suraswadi et al., 2003). Most mangroves are, however, relatively shallow, and a significant re-aeration take places with the atmosphere promoting good oxygen conditions in the mangrove waterways for large grazing organisms (Suraswadi et al., 2003). Loading of the mangrove forest with oxygen consuming substances, e.g. released from shrimp farms may, however, give rise to problems with low oxygen tensions (Burford & Longmore, 2001).

1.4

Organic matter sources in mangrove forests

The detritus in mangrove ecosystems is either derived from autochthonous (e.g. mangrove litter, phytoplankton, benthic algae) or allochthonous sources (e.g. river run-off, antropogenic discharges), and the fate of this detritus is recycling within or outwelling from the mangrove forest in dissolved microparticulate and macro-particulate (litter) forms (Chansang & Poovachiranon, 1990; Wattayakorn et al., 1990; Ayukai et al., 1998; Kristensen et al., 2000; Wattayakorn et al., 2000). The magnitude of litter fall, which is one of the most important autochthonous sources, depends on geographical locations, seasons and tree species (Twilley et al., 1986; Woodrooffe et al., 1988; Bunt, 1995; Shunula & Whittick, 1999). The fallen litter is potentially exported from mangrove forests by tidal exchange. This process is largely influenced by mangrove vegetation, where dense trees in the mangrove stand slow down water movements, favouring trapping within the forest (Furukawa & Wolanski, 1996; Furukawa et al., 1997). Chansang and Poovachiranon (1990) emphasized the role of litter trapping processes inside a mangrove forest in Thailand, as only 0.5–15% of litter fall was exported. Moreover, feeding and burial activities of mangrove crabs can remove a substantial amount (28–100%) of leaves before litter is exported to coastal waters (Robertson, 1986; Poovachiranon & Tantichodok, 1991). Nevertheless, export of mangrove litter is an important process for nutrient transports between mangrove ecosystem and outer boundary (Chansang & Poovachiranon, 1990; Wattayakorn et al., 1990). Only little information is available on the dynamics of the autochthonous sources (Poovachiranon et al., 2003). Litter fall was collected for a two year period along the main creek of a mixed mangrove forest in Thailand. While litter from only six mangrove species were observed within the litter fall traps, samples of tidal litter transport revealed a total of 23 mangrove species. Rhizophora apiculata and Ceriops tagal, followed by R. mucronata and Xylocarpus granatum were the most dominant species. The litter fall ranged from 8 to 12 tons ha−1 yr−1, and was mainly composed of leaves (78%),

8

BIOGEOCHEMISTRY OF MARINE SYSTEMS

followed by apexes, twigs and fruits (Fig. 1.3). Spatial variations in litter fall were prominent and indicated changes in mangrove and environmental conditions. While seasonal variation in leaf fall was negligible, fruit and flower fall occurred during confined periods. Year-to-year variations in litter fall, on the other hand, were prominent, with a higher rate recorded in dry than wet years. Less than 1% of total litter fall was exported from the mangrove forest, of which the net exports were highly correlated with spring tides. Twigs and fruits were major fractions of the exported litter, while leaves only comprised 28% of the total export (Fig. 1.3). Most litter was retained within the mangrove forest due to trapping effects of the dense mangrove vegetation, consumption and burial by sesarmide crabs. This mangrove forest had a high nutrient retention capacity, as less than 1% of nitrogen and phosphorus in fallen litter were directly exported out of the mangrove via litter materials. The remaining portions were efficiently recycled in the mangrove forest and subsequently outwelled from the mangrove in dissolved and micro-particulate forms. Other autochthonous sources are phytoplankton, benthic algae and macrophyte production, which are considered to be important in large lagoons with fringing mangrove forests (Robertson et al., 1992). Where shading is not severe, prop root epiphytes may be highly productive, and high benthic microalgae production has been found to be significant on creek banks and mud flats with high radiation (Kristensen et al., 1988). 100 Litter fall Litter transport

Litter fall and transport (%)

80

60

40

20

0 Leaf

Twig

Fruit

Apex

Misc.

Fig. 1.3 Annual litter fall and litter transport in the Bangrong mangrove forest, Phuket, Thailand. The litter is separated into leaves, apexes, twigs, fruits and miscellaneous (misc., unidentified) material. Modified from Poovachiranon et al. (2003).

MANGROVES OF SOUTHEAST ASIA

9

The identification and quantification of allochthonous sources are even less documented. The increasing number of shrimp farms in Southeast Asia may contribute significantly to inputs of dissolved nutrients as well as particulate nutrients during flushing of the shrimp farms (Burford & Longmore, 2001). River-influenced mangroves may be affected by allochthonous sources, e.g. land run-off, agriculture and sewage discharges, but very limited information is available. Urbanization is rapid in Southeast Asia, and it is very likely that organic matter sources will change during the coming decades. The majority of mangrove forests, however, occur in estuarine areas, or as dense forests with tidal channels. These mangrove systems have a high productivity, and the carbon fixed by mangrove vegetation is likely to be the dominant contributor to detritus (Lee, 1999). 1.4.1

Decomposition of detritus

During the last decades, there have been a number of studies on the degradation of mangrove litter by use of litter-bags (Wafar et al., 1997; Ashton et al., 1999). These have shown rapid initial decomposition followed by low rates and accumulation of refractory compounds. The detritus is often enriched by nitrogen during the decomposition (Wafar et al., 1997; Holmer & Olsen, 2002), and recent results also show enrichment with phosphorus (Nielsen & Andersen, 2003). The detritus derived from mangrove forests is generally low in nutrients and high in structural carbohydrates, aromatic humic compounds and geopolymers (Benner & Hodson, 1985; Kristensen et al., 1995). Nitrogen and phosphorus of mangrove leaves are re-absorbed and translocated before abscission (Woodroffe et al., 1988; Feller et al., 1999), and it has, for example, been found that the phosphorus concentration in the yellow leaves of Rhizophora sp. is only half of the content in fresh green leaves (Lugo et al., 1990). Litter fall of the evergreen mangrove trees occurs throughout the year, but often with a peak immediately before or during the wet season in monsoonal areas. The amount of leaves decomposing in and on the forest floor is a function of input (litter fall and import from adjacent areas) and output (export by tides and runoff, decomposition and removal by leaf-eating crabs). The residence times of detritus before export vary between 58–252 days in basin mangrove forests and 40–91 days for fringe and riverine forests (Twilley et al., 1986). The leaves on the forests floor range from yellow to black in colour, indicating the decomposition. Decomposition rates increase with humidity, temperature and oxygen availability, and depend on the composition of the organic matter. Often litter accumulates in the higher and less frequently inundated parts of the forest floor, while a smaller amount has been found in the lower zones of Southeast-Asian fringe forests. The rate of decomposition of mangrove leaves varies among species. Avicennia leaves decompose at a faster rate than those of other species,

10

BIOGEOCHEMISTRY OF MARINE SYSTEMS

probably because the leaves are thinner and sink immediately to the sediment surface, compared to the thicker and buoyant leaves of e.g. Sonneratia and Rhizophora (Wafar et al., 1997). The loss of organic matter is generally rapid during the first couple of weeks due to leaching of soluble organics (Fig. 1.4). The increase in nitrogen content, by 2–3 times, during decomposition is most likely associated with the accumulation of nitrogenous biomass of microorganisms colonizing the decaying litter (Fig. 1.4). Nitrogen fixation is generally low in mangrove sediments, diminishing the potential for this process to account for the nitrogen enrichment of the detritus (Kristensen, 1997), whereas relatively large nutrient pools provide a potential for bacterial nitrogen incorporation directly from pore water pools (Kristensen et al., 1998; Holmer et al., 2001). 120 Leaf litter remaining (% of initial AFDW)

100 80 60 40 20 0

0

10

20

30

40

50

60

70

80

90

% nutrient remaining

200

150

100

50 Nitrogen Phosphorus

0 0

10

20

30

40

50

60

70

80

90

Time (days) Fig. 1.4 Weight loss and nutrient content in the remaining litter (yellow leaves of Rhizophora apiculata) during a decomposition experiment in the Bangrong mangrove forest, Phuket, Thailand. The leaves were buried in surface sediment in the mangrove in litter boxes and sampled with intervals during a 3 month period. Modified from Nielsen and Andersen (2002) and Holmer and Olsen (2002).

MANGROVES OF SOUTHEAST ASIA

11

The decay pattern of mangrove leaves shows rapid initial losses in biomass followed by a slower decrease, and a single exponential model has frequently been used to describe the decomposition process (Kristensen, 1997). The decomposition constants are generally low compared to marine detritus, and half-life ranges between 20 and 226 days depending on the type of detritus and position along the tidal gradient (Wafar et al., 1997; Kristensen, 1997; Ashton et al., 1999; de Boer et al., 2000; Holmer & Olsen, 2002). The decomposition is faster at the sediment surface compared to burial in the sediments, probably because water soaking causes leaching of labile materials and promotes leaf conditioning by microbes (Chale, 1993). The decomposition may be further enhanced by the presence of herbivorous crabs (Kristensen & Pilgaard, 2002; Thongtham et al., 2003). A study of the decomposition of R. apiculata leaves buried in surface sediments showed that phosphorus also accumulates with time in the decomposing material (Fig. 1.4, Nielsen & Andersen, 2003). The incorporation of phosphorus into the decomposing leaves was probably associated with binding to humic acids and metals, especially iron, which also accumulated in considerable amounts in the leaves. The phosphorus most likely originated from iron-bound phosphate in the sediment as the decomposition of buried leaf litter increased the sediment microbial respiration, which led to reduced redox potential and higher dissolved reactive phosphate concentrations in the sediment pore water. This binding of phosphate to refractory organic material and the presence of oxidized iron at the sediment–water interface resulted in low release of dissolved reactive phosphate from the sediment and thus enhanced the retention of phosphorus in the mangrove forest. Similar observations have been found for nitrogen with reduced fluxes and enhanced retention in the sediments (Holmer & Olsen, 2002). The decomposition of seagrass leaves in the mangrove, e.g. imported from the seagrass beds in front of the mangrove forests during high tide, is much more rapid than mangrove leaves. The seagrass material was degraded within the first week of the decomposition, probably because seagrass detritus is less refractory and has higher nutrient contents compared to mangrove leaves (Holmer & Olsen, 2002). Seagrass detritus may thus be an important organic matter source for the sediment microbial community and increase nutrient recycling in the mangrove forest. About 32–36% of the mineralized nitrogen from the seagrass leaves was released across the sediment–water interface, where it became available for primary producers or exported to adjacent ecosystems.

1.5

Sediment biogeochemistry

Sediment biogeochemistry in mangrove forests is just as complex as water column biogeochemistry. The sediments are highly heterogenic and the

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

composition varies from muddy sediments in the high-intertidal forest to coarse sands at the flats in front of the mangrove forest (Kristensen et al., 1995, 1998; Alongi et al., 1998). The sediment structure is considered to have a major impact on the abiotic processes, e.g. diffusive and advective processes, and the biotic conditions, e.g. species diversity and abundance, which among other things play a crucial role for the recycling of organic matter and the oxidation of the sediments. The sediments are characterized by the abundance of silt and sand with minor amounts of clay (Badarudeen et al., 1996). Some mangrove sediments show enhanced contents of calcium carbonate as a result of the large quantity of shell fragments in the sediments (Badarudeen et al., 1996). The supply of organic matter to the sediments varies as described above, but the significant transportation of detritus in the mangrove creates a large potential for organic matter inputs to the sediments. The organic matter is primarily introduced through sedimentation, but bioturbating organisms may also influence the organic matter cycling. The sesarmine crabs, for example, crawl up into the forest trees, collect leaves and use these to cover their burrows. This habit adds organic matter to the sediments (Thongtham et al., 2003). The high- and midintertidal sediments are often densely vegetated with mangrove trees, and this may also provide a direct source of organic matter to the sediment, e.g. through root exudation and decay. Due to the large input and the refractory composition of the organic matter, mangrove sediments are often found to be organic-rich (Fig. 1.5), but as the detritus is relatively nutrient-poor and refractory, the sediments are generally characterized by low net mineralization rates (Kristensen et al., 1992, 1995). 1.5.1

Total microbial activity in mangrove sediments

Oxygen uptake and carbon dioxide production measured as fluxes across the sediment–water interface represent estimates of total microbial activity in the sediments. Sediment oxygen uptake is generally considered to be an integrated measure of aerobic bacterial respiration as well as chemical and chemoautotrophic oxidation of reduced metabolites (e.g. sulphide and ammonium). The release of carbon dioxide (CO2) from sediments, on the other hand, may provide a reasonable estimate of total benthic carbon mineralization when no carbonate precipitation or dissolution occurs. Only very few studies of sediment oxygen uptake and CO2 production have been undertaken in Southeast Asian mangrove forests. The benthic oxygen uptake rates in tropical mangrove forests are generally low compared to other intertidal environments at the same temperature (Kristensen et al., 1995, 1998). This is considered to be due to low microbial activity, controlled by the refractory and nutrient-poor sediment organic matter, which limits the decomposition potential (Kristensen et al., 1995). In mangrove forests, the lowest rates have been found for inundated sediments,

13

MANGROVES OF SOUTHEAST ASIA

5

0.3

0.2 3

POC

2 0.1 1

PON (%DW)

POC (%DW)

4

PON

0

High-intertidal Mid-intertidal Creek bank Low-intertidal

0.0

0.06

TP (%DW)

0.05 0.04 0.03 0.02 0.01 0.06

C : N, C : P (molar)

High-intertidal Mid-intertidal Creek bank Low-intertidal

100 75

C:P

50 25

C:N

0 High-intertidal Mid-intertidal Creek bank Low-intertidal

Fig. 1.5 Changes in sediment organic matter contents from high-intertidal mangrove forest to low-intertidal mudflats in the Bangrong mangrove forest, Phuket, Thailand. Upper panel: particulate organic carbon (POC) and nitrogen (PON), mid panel: total phosphorus content (TP) and lower panel: molar ratios between POC and PON and POC and TP. Modified from Holmer et al. (2001).

whereas air-exposed sediments generally have higher oxygen demand (Fig. 1.6). Up to 6 times higher oxygen uptake by air-exposed sediment relative to submerged sediment has been measured, and is considered to be caused by an increased area of oxic–anoxic interfaces combined with a reduced thickness of the diffusive boundary layer during air exposure (Kristensen et al., 1992).

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

3.0

O2 uptake submerged

2.5 2.0

Dry

1.5 1.0

wet

0.5 0.0 High-intertidal

Mid-intertidal

Creek bank

Dark flux (mmol m–2 d–1)

3.0

Low-intertidal

O2 uptake air-exposed

2.5 2.0 1.5 1.0 0.5 0.0 High-intertidal

Mid-intertidal

Creek bank

Low-intertidal

6 CO2 production submerged

5 4 3 2 1 0 High-intertidal

Mid-intertidal

Creek bank

Low-intertidal

Fig. 1.6 Changes in sediment oxygen uptake and CO2 production from high-intertidal mangrove forest to low-intertidal mudflats in the Bangrong mangrove forest, Phuket, Thailand. Oxygen uptake was measured in the wet and dry season and under submerged (upper panel) and air-exposed conditions (mid panel). The CO2 production was only measured under submerged conditions (lower panel). Modified from Holmer et al. (1999).

MANGROVES OF SOUTHEAST ASIA

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Drainage of water from sediment interstices (e.g. burrows and cracks) during low tide exposes sites that are otherwise anoxic to oxygen, thereby increasing the area of interfaces where rapid oxygen uptake can occur. The flux of carbon dioxide from inundated mangrove sediments is generally higher than oxygen uptake (Fig. 1.6), providing community respiratory quotients (CRQ) of 0.8–3.2 (Kristensen et al., 1991, 1992, 1994, 1995; Alongi et al., 1998; Holmer et al., 2001). CRQ values higher than 1 indicate that only a fraction of the sulphide produced by sulphate reduction is reoxidized by oxygen and also suggests that anaerobic mineralization processes are important for the total mineralization of organic matter in the sediments. The oxygen uptake and carbon dioxide production generally decrease with the degree of submergence (Fig. 1.6). Rates are highest in the high-intertidal organic-rich sediments and lowest in the mud and sand flats in front of the mangrove forests (Kristensen et al., 1995; Holmer et al., 2001). This is probably due to the large changes in physical and biological parameters along this gradient. The high-intertidal sediments are densely vegetated and have an active population of burrowing crabs, which ensures aeration of the sediments. Aerobic conditions promote the decomposition of refractory organic matter (Kristensen & Holmer, 2001). With increasing submergence, the organic matter content and the bioturbating activity decrease, and the sediments become more reduced and the microbial activity also decreases. These changes in total microbial activity along the transect from land to ocean have large effects on the nutrient regeneration within mangrove forests and will be discussed below (Holmer et al., 2001). 1.5.2

Mineralization pathways in mangrove sediments

Biogeochemical studies conducted in Thailand (Kristensen et al., 1991, 1995; Alongi et al., 1998) and Jamaica (Nedwell et al., 1994) suggest that aerobic respiration and sulphate reduction are major pathways of organic matter diagenesis in mangrove sediments, but recent studies have shown that iron reduction may also be an important respiration process (Alongi et al., 1998; Kristensen et al., 2000). Generally, the mineralization of organic matter in mangrove sediments occurs through the pathways found in marine sediments (Kristensen et al., 1995), and thus takes place with a number of electron acceptors (Fig. 1.7). Oxygen is an important and efficient electron acceptor, but is rapidly consumed and is usually only present in the upper millimetres of temperate coastal sediments (Revsbech et al., 1986). About 50% of the organic matter oxidation is considered to take place under aerobic conditions (Canfield, 1993). In the deeper layers, electron acceptors such as nitrate, iron and sulphate become important for the mineralization. The oxygen penetration into mangrove sediments has been found to be quite limited (Andersen & Kristensen, 1988), although the penetration depth may

16

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Water

Burial:

Fig. 1.7 Mineralization pathways and products in marine sediments. The width of the arrows is a rough indicator of the importance of the individual fluxes and reactions in continental margin sediments. Modified from Thamdrup and Canfield (2000).

have been underestimated in the permeable coarse sediments of the mangrove creeks and flats in front of the mangrove forests due to methodological limitations. The oxygen penetration depth is similar to subtidal temperate coastal sediments in the order of 1–5 mm (Revsbech et al., 1986). Despite this limited penetration, the oxygen respiration appears to be very important for the organic matter decomposition in mangrove sediments (Kristensen, 1997). This is probably due to a combination of the efficiency of the electron acceptor and the presence of labile organic matter in the surface sediments. A general characteristic of mangrove ecosystems is low levels of both particulate and dissolved (Fig. 1.5) nitrogen compounds (Fig. 1.8), suggesting that the biogeochemical cycling of nitrogenous compounds is closely coupled between assimilation and dissimilation processes. In particular, the concentration of nitrate is low in tropical sediments, and denitrification is considered to be of minor importance for the decomposition of organic matter (Kristensen et al., 1998). The primary cause for nitrogen limitation is the low nitrogen content of mangrove litter, and it has been suggested that mangrove forests are efficient in retaining and recycling nitrogen via several mechanisms that reduce export (Alongi et al., 1992). These mechanisms include re-absorption or re-translocation of nitrogen prior to leaf fall, burial of fallen detritus by crabs and rapid and efficient uptake of dissolved materials by bacteria. Nitrogen mineralization is often counterbalanced by ammonium assimilation, and the low net mineralization rates result in low pore water pools of ammonium and low release from the sediments (Fig. 1.9). The concentrations are, in particular, low in vegetated zones, where nutrients are taken up by the vegetation through the roots

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+

NH4 (mM) 0

50

100

150

200

250

300

0 High-intertidal Mid-intertidal Creek bank Low-intertidal

5

10

Depth (cm)

15

20 3–

PO4 0

4

(mM) 8

12

0

5

10

15

20 Fig. 1.8 Pore water concentrations of ammonium (upper panel, NH4+) and phosphate (lower panel, PO43−) in the Bangrong mangrove forest, Phuket, Thailand during the dry season. Nitrate concentrations were <1 μM. Modified from Holmer et al. (2001).

(Kristensen et al., 1998). Although the waters flooding mangrove sediments generally are low in ammonium, there is rarely ammonium release from the sediment (Holmer et al., 2001). Ammonium fluxes are usually lower than and opposite in direction to those observed in temperate subtidal and intertidal areas. Since the profiles of pore water ammonium indicate an upward diffusive flux, rapid consumption of ammonium must occur near the sediment–water interface. The most likely agents for ammonium removal in the uppermost sediment layer are assimilation by benthic microalgae and nitrification (Kristensen et al., 1998; Holmer et al., 2001), but immobilization by an acive, litter-degrading aerobic bacterial community may also contribute significantly (Holmer & Olsen, 2002).

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

Nutrient flux (mmol m–2 d–1)

0.0

–0.2 –0.4 –0.6 –0.8 –1.0

Nitrate Ammonium Phosphate High-intertidal

Mid-intertidal

Creek bank

Low-intertidal

Fig. 1.9 Changes in nutrient fluxes (nitrate, ammonium and phosphate) across the sediment–water interface from high-intertidal mangrove forest to low-intertidal mudflats in the Bangrong mangrove forest, Phuket, Thailand. Modified from Holmer et al. (2001). Negative values represent uptake.

Dynamics and fluxes of nitrate in marine sediments are primarily driven by two key processes, nitrification and denitrification. Also, the recently discovered anaerobic oxidation of ammonium coupled to nitrate reduction, producing N2, may significantly influence the fate of nitrogen compounds (Thamdrup & Dalsgaard, 2002). The process is known as anammox and is particularly important in low activity environments, but remains to be examined in mangrove sediments. Low pore water concentrations of nitrate in mangrove sediments are probably a consequence of low nitrification rates combined with high nitrate consumption by the sediment community. Potential rates of sediment nitrification in mangrove sediments are lower than usually found in temperate coastal sediments (Kristensen et al., 1998). Although the gradient of nitrate across the sediment–water interface in mangrove forests indicates a diffusive release into the overlying water, measured fluxes are generally directed into the sediment (Fig. 1.9). Recent studies have shown that measured rates of denitrification can only account for 10–30% of the nitrate uptake by the sediment, whereas the rest is supplied from coupled nitrification–denitrification in the sediments (Kristensen et al., 1998). The fate of nitrate at the sediment–water interface in mangrove sediments is considered to be similar to ammonium: consumption by an active community of benthic microalgae or incorporation into decomposing bacteria (Holmer et al., 2001; Holmer & Olsen, 2002). These processes serve as efficient mechanisms for nutrient conservation. The availability of sediment nutrients to microbes and plants is complicated by geochemical processes, such as the involvement of some nutrients in

MANGROVES OF SOUTHEAST ASIA

19

adsorption reactions to clay minerals. The ammonium adsorption is low in mangrove sediments compared to temperate salt marsh sediments, probably due to higher concentration of competitive cations such as iron (Holmboe & Kristensen, 2003). A study of mangrove forest sediments in Thailand has shown that the adsorption of ammonium is negatively related to the organic content of the sediments, which is in contrast to findings for other marine sediments (Holmboe & Kristensen, 2003). It was suggested that organic material may have a diluting effect on the exchange capacity in fine-grained sediments, and that organic coatings may block ion exchange sites on clay surfaces. Ammonium availability was thus found to be relatively higher compared to temperate sediments. Only very few studies have examined iron reduction in mangrove forest sediments, but some have been conducted in the Southeast Asian region. In subtidal sediments, rates of iron reduction have found to be positively correlated with the pool of oxidized iron (Thamdrup, 2000), and this relationship may be useful in predicting the potential of iron reduction in mangrove forest sediments. The presence of oxidized iron is considered to be controlled by the origin and the oxidation of the sediments. It can thus be expected that iron reduction will be most important in the high- and mid-intertidal sediments with a large contribution of terrigenous compounds, many bioturbating organisms and oxidation of the sediments by vegetation. This is consistent with a study of a mangrove forest in Thailand (Kristensen et al., 2000), where it was found that iron reduction accounted for up to 70–80% of the total carbon oxidation in mid-intertidal rooted mangrove forest sediments (Fig. 1.10). Similarly, Alongi et al. (1998) found that iron reduction was most important in a regenerating mangrove forest, where the iron pools were at maximum compared to older forests. The contribution from iron reduction was less in the low-intertidal sediments (30–40%) and almost absent in the seagrass sediments in front of the mangrove (<1–15%) (Fig. 1.10, Kristensen et al., 2000). The shift in dominance of electron acceptors along the tidal transect was found to be related to the presence of roots and benthic fauna, but also the sediment composition (grain size, organic content and iron content) appeared to be an important co-factor. The grain size may be an important controlling factor for the iron content, as iron oxides are adsorbed to sediment surfaces, and the decrease in iron content along the transect was coincident with an increase in the sediment particle size (Kristensen et al., 2000). There has been relatively more focus on measurements of sulphate reduction in mangrove sediments, mainly since sulphur cycling has been intensively studied during the last three decades, but also because the sulphur cycling in mangrove sediments can have significant impacts on the benthic community due to a variety of secondary effects, e.g. associated pH changes. Sulphate reduction is not expected to be controlled by the concentration of sulphate in the mangrove forest sediments due to the high salinities generally found here.

20

BIOGEOCHEMISTRY OF MARINE SYSTEMS

Total CO2 production (mmol C cm–3 d–1)

Depth (cm)

0.0 0

0.5

1.0

1.5 0.0 0

5

5

10

10

15

Mid-intertidal 15

0.0 0

0.5

1.0

1.5 0.0 0

5

5

10

10

15

Low-intertidal 15

0.5

1.0

1.5

Creek bank 0.5

1.0

1.5

Seagrass bed

Fig. 1.10 Iron reduction and total CO2 production obtained by incubation of sediment in closed jars at four sites in the Bangrong mangrove forest, Phuket, Thailand. The dark shaded area indicates the contribution from iron reduction. Modified from Kristensen et al. (2000).

It is much more likely that the rates are controlled by the availability of organic matter and the biological and physical processes acting on the oxidation of the sediments, e.g. bioturbation, root oxidation and tides. Sulphate reduction appears to be an important process in mangrove sediments, and relatively high rates have been found (e.g. Kristensen et al., 1995; Kristensen, 1997; Alongi et al., 1998; Holmer et al., 1999). This suggests that sulphate reduction may contribute significantly to mineralization of organic carbon and nutrient availability in

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21

tropical mangrove sediments. Patterns in sulphate reduction rates have been found to reflect differences in physical and biological conditions in a Thailand mangrove forest (Holmer et al., 1994). Highest rates were found at the vegetated site within the mangrove (R. apiculata) forest, probably due to high input of labile organic matter from sedimentation of phytoplanktonic detritus among the stilt roots or by root exudates and root decay (Fig. 1.11). The contribution of sulphate reduction to CO2 production varies among studies from about 20% up to 85%, most likely due to the heterogeneity of these sediments (Holmer et al., 1994, 1999). Sulphate reduction rates were at minimum at the highintertidal site, and this site was characterized by a large number of crab burrows deep into the sediments and dry conditions, as the sediments were flooded only for short periods of time during high tide (Fig. 1.11). These sediments were thus highly oxidized as revealed by positive redox potentials favouring mineralization processes that employ more oxidized electron acceptors than sulphate. Sulphate reduction accounted for 11% of the CO2 production. The rates were in between these two extremes at the mud flats in the mangrove waterways, whereas they increased again in the seagrass sediments in front of the mangrove, probably due to input of fresh organic matter from seagrasses or microphytobenthos, and sulphate reduction accounted for 20–40% of the mineralization (Holmer et al., 1999). The accumulation of sulphur in subtidal marine sediments is primarily controlled by the rate of sulphate reduction and the oxidation state of the sediment (Thode-Andersen & Jørgensen, 1989). In tidal environments, however, additional factors must be considered. Tidal currents and wave action can affect the oxidation status of sediments directly by increased advective transport of pore water and particles. During low tide, the sediment surface desiccates and oxygen can penetrate deeper into the sediment via burrows and cracks in the surface. The presence of rooted vegetation also strongly affects the biogeochemical cycling of sulphur by vertical translocation of organic matter and oxygen (Holmer & Nielsen, 1997; Holmer & Laursen, 2002), and the cycling of sulphur is closely coupled to the reactive iron pools (Thamdrup, 2000). Reactive iron oxides present in sediments may efficiently oxidize reduced sulphides. This suggests that the cycling of reduced sulphur compounds is highly dynamic in mangrove forest sediments. Buried pools of sulphides are usually extracted in two major pools, the acid volatile (AVS) and chromium reducible (CRS) pools. The AVS pools consist primarily of dissolved H2S and FeS compounds, and the CRS pools are mostly pyrite. AVS is considered to be the initial product formed, but the pyrite formation is also rapid, in particular in relatively oxidized sediments, such as vegetated mangrove forest sediments (Holmer et al., 1994). Pyrite appears to be the most important inorganic sulphur component in mangrove sediments, attaining pool sizes 50–100 times higher than acid volatile pools (Kristensen et al., 1991, 1992; Holmer et al., 1994, 1999; Alongi

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SRR (mmol m–2 d–1)

25

20

Wet

15

10

5 Dry 0 High-intertidal

Mid-intertidal

Creek bank

Low-intertidal

Seagrass bed

0 High-intertidal

Mid-intertidal

Creek bank

Low-intertidal

Seagrass bed

Mid-intertidal

Creek bank

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Seagrass bed

% SR of CO2 production

50

40

30

20

10

TRS pool (mmol S m)

40

30

20

10

0 High-intertidal

Fig. 1.11 Depth integrated rates of sulphate reduction (upper panel, SRR) and the contribution of sulphate reduction to total sediment metabolism measured as CO2 production across the sediment–water interface (mid panel, %SR of CO2 production) in the Bangrong mangrove forest, Phuket, Thailand. The depth integrated pools of total reducible sulphides are given in the lower panel (TRS pool). Results from dry and wet seasons are given. Modified from Holmer et al. (1999).

MANGROVES OF SOUTHEAST ASIA

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et al., 1998). There is quite a significant burial of sulphides in the mangrove forest sediments (Fig. 1.11), in particular in the mid-intertidal sediments, where the oxidation by bioturbating organisms is low and the sulphate reduction activity high, but also the high-intertidal sediments show large accumulations below depths of bioturbation (Holmer et al., 1994). This may give rise to major problems during land reclamation, as it is often these vegetated areas which are utilized for new land or shrimp farms. The exposure of reduced sediments with high sulphur pools to oxygen produce sulphuric acid during reoxidation, and thus creates highly acidic environments not suitable for shrimp farming or agriculture (Kryger & Lee, 1995; Kautsky et al., 2000). The burial of sulphides is much less in the non-vegetated sediments of mangrove creeks and in the mud flats, probably caused by oxidation facilitated by advection during shifting tides and air exposure during low tides. Also, here pyrite is the most important sulphur pool. There are, however, large site differences in the burial of sulphides in mangrove forests. The sulphur pools were low in the Bangrong mangrove forest, Thailand, compared to other mangrove sites, e.g. the nearby Ao Nam Bor mangrove. Similarly, Alongi et al. (1998) found that the pyrite concentrations increased with forest stand age suggesting that the sulphur accumulates with time in the mangrove sediments. In addition to the burial of inorganic sulphur compounds, an accumulation of organic sulphur has been found in the deep sediments in the inner part of mangrove forests (Holmer et al., 1994). A similar accumulation of organic sulphur has been observed in mangrove peats (Altschuler et al., 1983), but the underlying mechanisms behind this accumulation are not well understood. However, the burial of inorganic sulphur appears to be limited by the availability of iron (Holmer et al., 1994), which may favour formation of organic sulphur compounds. As sulphate has been found to be present to deep depths in the mangrove sediments examined so far (e.g. Holmer et al., 1994; Alongi et al., 1998), methane production is not considered to be an important mineralization process in mangrove forests. Methane-producing bacteria are able to compete only with sulphate reducers for electron donors at low sulphate concentrations (<200 μM), and such low concentrations have not been measured at depths of 30–50 cm examined so far. The sulphate concentration may still be as high as 8 mM at these depths (Holmer et al., 1994; Alongi et al., 1998). 1.5.3

Phosphorus cycling

Unlike nitrogen, iron and sulphate transformations, which are influenced by biological processes, chemical processes are more important for phosphorus transformations in aquatic environments. Phosphorus is an important element in mangrove forests, as the mangrove trees and other mangrove primary producers as well as the bacteria are often found to be phosphorus limited (Boto & Wellington,

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1983; Alongi et al., 1992; Feller et al., 1999). The phosphorus dynamics in mangrove sediments are considered to be closely coupled to the activity of mineralizing iron- and sulphate-reducing bacteria. Phosphate is largely insoluble and regarded as non-reactive in the presence of oxidized iron, but is easily solubilized via chemical or biological reduction of iron under anoxic conditions. Moreover, a large fraction of the apparently reactive pool of phosphorus is prone to be immobilized by strong adsorption to clay particles (Alongi et al., 1992). Thus, phosphate is generally found in very low concentrations in mangrove waterways where suspended iron-rich clay particles are abundant (Wattayakorn et al., 1990; Trott & Alongi, 1999; Ayakai et al., 2000), and is usually also very low in the pore waters (Fig. 1.8) where it is adsorbed to the sediment particles (Tam & Wong, 1995; Alongi et al., 1998; Holmboe et al., 2001). The flux of phosphate across the sediment–water interface is usually very low and rate measurements are constrained by large variability due to the heterogeneity of mangrove sediments and analytical precision near the detection limits (Holmer et al., 2001). In many cases, the fluxes are directed into the sediment (Fig. 1.9), and this high affinity of mangrove sediments has been ascribed to microbial uptake (Alongi et al., 1992), adsorption onto iron oxides (Holmboe et al., 2001) and binding to organic material (Hesse, 1963). The low phosphorus content of the detritus has been suggested to limit the potential for release of phosphate during mineralization (Fig. 1.5). At the same time most mangrove sediments show positive redox potentials in the surface sediments, favouring the binding of phosphate to oxidized iron. Iron-bound phosphate has been found in most mangrove sediments with highest concentrations at sites with high iron contents (Holmboe et al., 2001). In accordance with these observations, the pore water concentrations of phosphate have been found to be very low (<10 μM) in the surface sediments and increasing with depth (Alongi et al., 1998; Holmer et al., 2001). The increase with depth is considered to be due to lower diffusion to the water column and less binding of phosphate to oxidized iron. Enrichment of the sediments with organic matter (Nielsen & Andersen, 2003) or removal of vegetation (Alongi et al., 1998) has been shown to increase the pore water pools of phosphate. In both cases, this was coincident with increased mineralization of organic matter and reduced redox potentials, suggesting that the binding of phosphate to oxidized iron decreased (Alongi et al., 1998; Nielsen & Andersen, 2003). The total content of phosphorus in mangrove sediments is usually low (Fig. 1.5) compared to temperate sediments (Alongi et al., 1998; Holmboe et al., 2001; Holmer et al., 2001), though some high values have been found in fine-grained sediments (Badarudeen et al., 1998). The content is usually higher inside the mangrove forests due to the deposition of fine particles rich in metal and organic content compared to the sand flat sediments (Holmboe et al., 2001). The surface concentrations are usually higher due to deposition and diagenesis of organic matter holding phosphorus and metals, and adsorption of phosphate

MANGROVES OF SOUTHEAST ASIA

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to oxidized iron (Jensen & Thamdrup, 1993). In mangrove forest sediments, however, the total phosphorus content is quite constant or increases with depth (Alongi et al., 1998; Holmer et al., 2001). This is probably due to the high degree of bioturbation and high redox potentials deep in the sediments. Only very few studies have examined the inorganic forms of phosphorus in mangrove forest sediments, and the extraction schemes vary among these studies. In Southeast Asia, two detailed studies (Holmboe et al., 2001; Nielsen & Andersen, 2003) are available, and these have been undertaken in the Bangrong mangrove, Thailand with the extraction method described by Jensen and Thamdrup (1993). The pools of easily exchangeable phosphorus are low, in particular in the fine-grained mangrove forest sediments. This is reflected to some extent in the higher pore water concentrations of phosphate of sand flat compared to mangrove sediments. The iron-bound phosphorus has been measured as a quantitatively important pool in many mangrove sediments; (Fabre et al., 1999; Holmboe et al., 2001; Nielsen & Andersen, 2003) in the Bangrong, iron-bound phosphorus comprised about 20% of the total phosphorus in mangrove and 10% in the sand flat sediments. The higher concentrations in the mangrove were correlated to higher iron pools. Phosphorus adsorbed to clay, aluminium and organic surfaces in mangrove sediments were mostly relatively low (0.4–8%) compared to the other fractions, and spatial and temporal changes were difficult to explain. The phosphorus found in the humic acid fraction was much more important, in particular in the mangrove sediments (15–42%) compared to sand flat sediments (2–8%). This is consistent with other findings of high proportions of phosphorus found in organic forms in mangrove sediments (Hesse, 1963; Alongi et al., 1992). The opposite pattern is found for calcium-bound phosphorus, which is lower in mangrove compared to sand flat sediments due to the increase in CaCO3 along this transect. The organic fraction of phosphorus was high and relatively similar among the stations (about 17% of total P), and is consistent with other findings (27–87% of total P) (Hesse, 1962; Fabre et al., 1999). 1.6 1.6.1

Factors influencing the biogeochemistry Effect of forest type and age

Rates of nutrient transformation and recycling are presumably also dependent on forest type and age, but only few studies have examined the influence of forest type (Alongi et al., 1993, 1998; Middelburg et al., 1996). Early studies have suggested that some edaphic characteristics are species-specific (Nickerson & Thibodeau, 1985), but subsequent studies have shown that apparent differences between forest types are more closely related to root density and frequency of tidal inundation than to species-specific ability of mangroves to affect sediment biogeochemistry and nutrient turnover (McKee et al., 1988). Nevertheless,

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differences in pore water chemistry and microbial activity among different forest types have been found. Alongi et al. (1993) studied three different forest types in Papua New Guinea: Rhizophora–Bruguiera, Nypa and Avicennia– Sonneratia forests. For most variables, variations within forest type were as great as, or greater than, differences between forest types. A high-intertidal Nypa site was most different in edaphic characteristics compared to five lowintertidal stations suggesting that differences among forest types are mainly a function of tidal elevation rather than species-specific ability of mangroves to influence redox and nutrient status. The decomposition of organic matter has been found not to be related to the forest stand age and age-related differences in rates of forest production. There were, however, differences in the dominance of specific diagenetic pathways in different forests (Alongi et al., 1998). Sulphate reduction accounted for most of the total benthic mineralization (75–125%) in the old forests (>15 years), whereas it constituted a much smaller proportion (42%) in the younger forest (2-year old). This is probably due to the higher concentrations of organic matter in the older forests. Rates of solute efflux across the sediment–water interface and vertical profiles of dissolved manganese, nitrite and nitrate suggest that the more oxidized pathways, such as manganese reduction and denitrification– nitrification, coupled with aerobic respiration, accounted for most oxidation in the young forest. 1.6.2

Influence of macrofauna

Herbivorous crabs are important in litter processing in the Indo-West Pacific region (Robertson, 1986; Lee, 1997), whereas the detritus food web in other forests, e.g. of the Caribbean region, is primarily driven by microbial processes (Odum & Heald, 1972). Litter consumption and burial by crabs may have profound impact on the organic matter flow and nutrient dynamics within mangrove forests and affect the exchange with the adjacent coastal zone (Robertson, 1986; Twilley et al., 1997). Manipulative experiments with herbivorous crabs have shown that crab bioturbation significantly decreased the level of ammonium and sulphide concentrations in mangrove sediments (Smith et al., 1991), which in turn may improve the productivity of the mangrove forest. Crab burrows also play an important role in affecting the groundwater flow in, and the chemistry of, otherwise compact mangrove sediments (Wolanski et al., 1992). Field studies of leaf removal by sesarmide crabs in mangrove forests in Thailand have shown that the crabs can remove about 75% of the total daily litter fall and green, yellow as well as brown leaves were consumed (Fig. 1.12). It was estimated that the total population of sesarmide crabs could consume 58% of the total leaf litter per year, in this case 1130 tons (Thongtham et al., 2003). A more detailed study of the economically important sesarmide crab

MANGROVES OF SOUTHEAST ASIA

27

Brown 55%

Yellow 17%

Green 29% Total consumption = 3.02 mg DW (g ww crab–1)d–1 Fig. 1.12 Proportional consumption of Rhizophora apiculata leaves (green, yellow and brown leaves) by the sesarmide crab Neoepisesarma versicolor. Modified from Thongtham et al. (2003).

Neoepisesarma versicolor was done, and it was found that the crab spends most of the time during the day in its burrow or resting by the entrance, and is mainly active during the night. It mainly feeds on mangrove leaves (62%) and scraped-off food materials from sediment surface (38%). The plant material mostly comprised small pieces of mangrove leaves. Supplementary laboratory studies of leaf consumption and leaf preferences on N. versicolor showed that the crab mostly feeds on brown leaf, followed by green and yellow leaves. A number of studies from Australia and Southeast Asia have documented that sesarmide crabs (Grapsidae) may consume 28–79% of the annual litter fall (Robertson, 1986; Robertson & Daniel, 1989). The remaining litter is either exported to adjacent waterways (Boto & Bunt, 1981; Chansang & Poovachiranon, 1990), decomposed aerobically at the sediment surface or anaerobically within the sediment (Robertson et al., 1992). The proportion of litter entering each pathway is largely determined by mangrove geomorphology, tidal flushing and type of litter (Robertson et al., 1992). Much of the litter handled by crabs will eventually enter the microbial food chain, either in the form of faecal material or as uneaten remains buried in the sediment (Giddins et al., 1986; Robertson, 1986; Lee, 1997). Intact leaves are, in contrast to faecal pellets, readily exported by tides before sinking, and thus lost from the mangrove ecosystem. Faecal pellets have been found to be more nutritious for the sediment microbes compared to refractory mangrove litter and sediment organic matter, and showed a 2 and 3–10 times faster rate of decomposition, respectively (Kristensen & Pilgaard, 2002). Introduction of sesarmine crabs (Uca sp.) into non-bioturbated mangrove sediments has been found to enhance the sediment organic matter decomposition (Nielsen et al., 2003b). The sulphate reduction activity was in particular

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enhanced. The introduction of crabs also affected the dynamics of the sulphur pools, and they were generally reduced despite stimulated sulphate reduction activity. This suggests that the crabs are able to oxidize the sediments through their construction of burrows. A direct effect from the burrow walls was, however, only found in sediments where both mangrove seedlings and crabs were introduced. Here the iron reduction became more important in the deeper layers, and the sulphate reduction activity decreased compared to the nonvegetated and non-bioturbated sediment. A detailed investigation of the sulphur and iron-dynamics in the surroundings of crab burrows showed that oxidative processes were very active in a zone of 5–10 mm from the burrow wall and the content of oxidized iron increased, whereas the content of reduced sulphides in the sediments was significantly reduced in this area. There are many other benthic organisms of importance in Southeast Asia mangrove forests, but detailed studies of their impacts on organic matter cycling and nutrient regeneration are lacking. Benthic fauna has in particular been found to stimulate the decomposition of refractory organic matter by introducing this organic matter to aerobic degradation either by ventilation of burrows or physical movement of organic matter from the deep sediment layers to the surface (Banta et al., 1999; Kristensen & Pilgaard, 2002). It is thus likely that benthic fauna in mangrove sediments play an important role for the regeneration of nutrients from the refractory organic matter in mangrove sediments, but this remains to be explored, and this research should be initiated before it is too late. A case study of the sesarmide crab N. versicolor shows that this species used to be very common in Thailand’s mangrove areas but its numbers have drastically decreased due to human consumption as well as other factors (Thongtham et al., 2003). Local people recall that N. versicolor was dominant in the Bangrong mangrove forest, but due to deforestation for shrimp farming, land development and especially over-exploitation for human consumption, its abundance has drastically decreased. As these leaf-eating sesarmide crabs have been found to play such a significant role in litter turnover and nutrients retention in mangrove ecosystems, their disappearance is of major concern. 1.6.3

Effect of seasonal variations on mangrove forest biogeochemistry

Some studies have observed seasonal variations in primary production and nutrient dynamics in the tropics, but the variation is often low due to the constant environment and the intense recycling of nutrients. Enhanced outwelling during rainy seasons may affect nutrient concentrations in mangrove forests stimulating primary production in the water column (Rivera-Monroy et al., 1998), and the coupling between mineralization and immobilization in the water column has been found to shift between the wet and the dry seasons (Suraswadi et al., 2003). Heavy rainfalls may also have major effects on water

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column characteristics. Thong et al. (1993) found that the concentrations of inorganic nitrogen in creek water increased by a factor of 10 after heavy rains. Organic nitrogen concentrations also increased under these conditions. The increase in inorganic nitrogen in the creek was greatest when the water drained off vast areas of the mangrove forest either: (a) after the flooding of high amplitude spring tides, or (b) after heavy rains. On some occasions, but not consistently, dissolved organic nitrogen and particulate nitrogen increased as water drained off the forest. In contrast, when there was no drainage due to rain and the waters were generally confined within the creek banks, either during neap or spring tides of low amplitude, the inorganic nitrogen concentrations remained relatively low. These relationships suggest that the floor of the mangrove forest is a major source of inorganic nitrogen to the creek. Export of increased nitrogen levels from the mangrove creek to the coastal waters seems likely to be greatest on occasions of high rainfall. In addition to the low seasonal variation in parameters such as temperature and light, which are the most important factors of seasonal variations in temperate regions, the large spatial variations and the lack of long time-series in measured sediment processes contribute to the poor understanding of seasonal variation in the tropics. Most studies so far have concluded that the seasonal variations are less important compared to physical and biological processes acting on the mangrove forest sediments (Kristensen et al., 1995; Alongi, 1996). Only in seagrass beds has enhanced anaerobic mineralization been found during the rainy season, probably due to lower light availability which decreases the oxidation of the sediments (Blackburn et al., 1994; Holmer et al., 2001).

1.7

Sediment biogeochemistry and implications for mangrove vegetation

The biological conditions in mangrove sediments range from aerobic through to anoxic and highly reduced. Mangrove plants growing in waterlogged soils may be adversely affected by either the strongly reduced conditions or the accumulation of soluble phytotoxins including reduced iron, manganese and organic gases (McKee, 1993). In addition, species growing in anaerobic marine sediments must also cope with toxic concentrations of sulphides (Allam & Hollis, 1972). In most wet land plants, the first line of defence against these toxic soluble ions is to render them insoluble at the root surface by oxidation with air that diffuses to the root from the photosynthesizing shoots (Armstrong et al., 1994). However, Youssef and Saenger (1996) have shown that, unlike mature vegetation, roots of mangrove seedling have a very limited capacity for oxidative detoxification of the rhizosphere under laboratory conditions. Furthermore, detoxification by direct oxidation is even more

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limited for seedlings under field conditions, because the seedlings are commonly partially or totally submerged or growing at low irradiances under the mature stand. Youssef and Saenger (1996) found that seedlings of the viviparous mangroves showed no significant responses to root anoxia, but sulphides inhibited photosynthetic gas exchange and caused complete stomatal closure in many species. Stomatal closure was probably the result of the damaging effect of sulphide ions on root cell membranes. It was concluded that high concentrations of sulphides in mangrove sediments may adversely affect both growth and survival of mangrove seedlings at low irradiances. Direct oxidation of sulphides by oxygen is possible around roots of mature Avicennia marina, as indicated by a study by Andersen and Kristensen (1988). They measured oxygen concentration in aerial roots (pneumatophores), horizontal cable roots and surrounding sediment of A. marina in a mangrove forest in Thailand. The O2 concentration inside pneumatophores was 63–88% of air saturation, whereas the cable roots showed a lower concentration (62–73%) indicating an O2 gradient from the emerging parts to the subsurface roots. The O2 concentration in the roots was highest in the outer part of the aerenchyma. There was an oxic zone of the sediment around roots, but it was thin (ca. 0.5 mm), and suggests that only little O2 is released from the roots or that the oxygen is rapidly consumed by aerobic mineralization or reoxidation of reduced compounds. Also, Rhizophora sp. has been found to transport air to the buried portions of their roots through lenticels on the aerial portions, but there is usually no free oxygen in the sediments. There is, however, less hydrogen sulphide in the sediments under R. mangle and Avicennia germinans compared to other species, but it appears to be related to site differences (Carlson et al., 1983; Nickerson & Thibodeau, 1985). A. germinans plants reoxidize the sediments, as the mean sulphide concentration beneath them is nearly six times lower than that in immediately adjacent areas (Nickerson & Thibodeau, 1985), whereas R. mangle is only found in areas with low mean sulphide concentration not significantly different from nearby unvegetated soil. The effect of hydrogen sulphide and metallic sulphides on mangrove vegetation during reforestation has only been briefly studied. Avicennia spp. with underground cable roots and pneumatophores, show clear signs of degeneration on mangrove plots, 12–14 years after reclamation of the land (Kryger & Lee, 1995). By contrast, saplings of Rhizophora spp. with aerial stilt roots were observed to establish within this area. It was found that concentrations of hydrogen sulphide in mangrove soils increased with the age of the soil, and it is suspected that the build-up of sulphides may affect the growth of different types of mangrove vegetation. The degeneration of Avicennia may be caused by the large deposition of silt, brought from inland construction areas through the inland drainage water, creating oxygen stress (Kryger & Lee, 1995). It is suspected that gaseous hydrogen sulphide may aggravate this condition of

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hypoxia, hence causing Avicennia spp. to die. This was supported by the observation that, at the mature plot where the gaseous hydrogen sulphide content was high, the population of Avicennia spp. with underground cable roots was low, and seedlings of these species were not able to survive well. Conversely, Rhizophora spp., which have aerial stilt roots that also help in the breathing processes, were abundant. Aerial roots could help the plants to overcome the high concentration of gaseous hydrogen sulphide in the soil (Kryger & Lee, 1995).

1.8

Biogeochemistry in mangroves affected by anthropogenic activities

A recent review of mangrove research concludes that there is a strong bias towards subject areas such as floristic and basic plant ecology rather than ecological processes and nutrient cycling in tropical Asia (Li & Lee, 1997). Due to the capacity of mangrove forests to withstand and retain excessive nutrients, there has been considerable focus on these environments as a natural treatment facility for effluents from, e.g. aquaculture (Robertson & Phillips, 1995; Sansanayuth et al., 1996; Wong et al., 1997; Tam & Wong, 1999), but while various potential impacts of activities such as prawn farming have been proposed (Primavera, 1993), the effects of not only mangrove removal but also the operation of the new use remain to be addressed. Robertson and Phillips (1995) attempted to evaluate the impact of shrimp pond effluent in relation to the assimilative capacity of mangroves for nutrients (N and P). It was estimated that between 2 and 22 ha of mangrove forests are required to strip the N and P in effluent from each hectare of prawn pond. These authors nevertheless discuss that the calculation probably cannot be directly extrapolated to larger scales and direct experimentation on the response of mangrove forests to pond effluent has to be conducted. A nutrient enrichments study shows that anoxic sediment decomposition was not affected by nutrient enrichments at the level applied (2–3 times background concentration) (Holmboe et al., 2001). This was substantiated by a low nutrient release from mangrove sediments and suggests a fast turnover of nitrogen and phosphorus by nutrient deficient bacteria in the mangrove sediments. Analysis of phosphorus fractions in the mangrove sediment revealed that added excess phosphate was efficiently taken up by the sediment particles and primarily retrieved in the easily exchangeable and ironbound fractions. The mangrove forest sediments thus acted as phophorus sinks. The long-term capacity of mangrove forests to retain excessive nutrient discharges is not known yet, but may be limited because most of the pollutants accumulate in the top layer (0–1.5 cm) of the sediment (Tam & Wong, 1999). In some cases, anthropogenic interference to mangrove forests may not always result in undesirable consequences. Li and Lee (1998) reviewed the particulate

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organic matter dynamics of a small, mangrove-line embayment in the Pearl River estuary in southern China and concluded that the recent increase in waterfowl numbers supported is probably related to increased food supply from domestic sewage discharge. Many countries in Southeast Asia have experienced almost an explosive growth in the number of shrimp farms and a corresponding reduction in the mangrove forests area (Platong, 1998; Kautsky et al., 2000). The increasing commercial shrimp farming in Thailand is considered to possess the most important threat on the stability of the coastal zone and the mangrove forests (Fig. 1.13, Platong, 1998). One of the largest problems associated with shrimp farming is the deposition and removal of the organic and nutrient rich sludge, which accumulate in the farms during the production cycle. The sludge deposits have traditionally been released through the mangrove waterways, but due to the fear of acute and long-term eutrophication problems, the recent legislation has forbidden this practice in favour of dry deposition on shore. As the onshore deposition has a high demand for resources, investigations have been undertaken to find alternative and cheaper solutions, e.g. controlled deposition where the mangrove acts as a biological filter of the enhanced nutrient loading (Robertson & Phillips, 1995; Wong et al., 1995; Sansanayuth et al., 1996; Gautier et al., 2001). Results from a study of shrimp farm impacts show that the benthic mineralization is doubled in the shrimp farm deposits compared

800 Shrimp production 700

250 Mangrove area

600

200 500 150

400 300

100

200

Mangrove area (103 ha)

Annual shrimp production (metric tons)

300

50 100 0

0 1978 1980

1982 1984 1986

1988 1990 1992 1994 1996

Fig. 1.13 Comparison between shrimp production (metric tons) and mangrove area (103 ha) in Thailand from 1979 to 1996. Modified from Platong (1998).

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0 months 7 months 16 months 22 months

CO2 release (mmol m–2 d–1)

150

100

50

0 Deposited shrimp pond waste

Mangrove sediment

Fig. 1.14 Changes in CO2 release (mmol m−2 d−1) from shrimp pond waste deposits and adjacent mangrove sediment in an upper mangrove area in the Ranong mangrove forest, Thailand. Modified from Nielsen et al. (2003a).

to untreated mangrove forests (Fig. 1.14). The enhanced mineralization was entirely due to sulphate reduction creating more reduced conditions in the sediments and possibly negative effects on the mangrove vegetation. The enhancement was, however, short-lived and after one year the carbon mineralization was already reduced to the untreated level. Introduction of burrowing crabs and mangrove seedlings in the shrimp farm deposits is essential as they improve sediment conditions in general, and accelerate reoxidation of reduced sulphide by introducing oxygen deep into otherwise anoxic sediment layers (Nielsen et al., 2003a). However, long time scales and deep and intense bioturbation are needed to observe major changes in sludge biogeochemistry, as immature vegetation and infaunal communities only have limited impact. Thus, deposition of shrimp pond sludge into mangrove areas can reduce some of environmental threats related to onshore deposition especially if deposited in areas with intense infaunal activity and high degree of vegetation. But again, the long-term environmental impacts of this practice remain unknown. Questions such as whether conversion of natural mangrove forests to shrimp ponds will lead to irreversible long-term ecological losses are still equivocal. Further research that is specifically directed towards answering practical challenges to our understanding of tropical mangrove forest ecology and biogeochemistry is urgently needed. Despite decades of increased research effort, answers to many fundamental questions concerning practical management issues of tropical mangrove forests are still largely unavailable.

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Kristensen, E., Holmer, M. and Bussawarit, N. (1991) Benthic metabolism and sulfate reduction in a south-east Asian mangrove svamp. Marine Ecology Progress Series, 73, 93–103. Kristensen, E., Holmer, M., Banta, G., Jensen, M.H. and Hansen, K. (1995) Carbon, nitrogen and sulfur cycling in sediments of the Ao Nam Bor mangrove forest, Phuket, Thailand: a review. Phuket Marine Biological Centre, 60, 37–64. Kristensen, E., Jensen, M.H., Banta, G.T., Hansen, K., Holmer, M. and King, G.M. (1998) Transformation and transport of inorganic nitrogen in sediments of a Southeast Asian mangrove forest. Aquatic Microbial Ecology, 15, 165–175. Kristensen, E., King, G.M., Holmer, M., Banta, G.T., Jensen, M.H., Hansen, K. and Bussarawit, N. (1994) Sulfate reduction, acetate turnover and carbon metabolism in sediments of the Ao Nam Bor mangrove, Phuket, Thailand. Marine Ecology Progress Series, 109, 245–255. Kryger, L. and Lee, S.K. (1995) Effects of soil ageing on the accumulation of hydrogen sulphide and metallic sulphides in mangrove areas in Singapore. Environment International, 21, 85–92. Lee, S.Y. (1997) Potential trophic importance of the faecal material of the mangrove sesarmine crab Sesarma messa. Marine Ecology Progress Series, 159, 274–283. Lee, S.Y. (1999) Tropical mangrove ecology: Physical and biotic factors influencing ecosystem structure and function. Australian Journal of Ecology, 24, 355–366. Li, M.S. and Lee, S.Y. (1997) Mangroves of China: a brief review. Forest Ecology and Management, 96, 241–259. Li, M.S. and Lee, S.Y. (1998) The particulate organic matter dynamics of Deep Bay, eastern Pearl River estuary, China I. Implications for waterfowl conservation. Marine Ecology Progress Series, 172, 73–87. Lugo, A.E., Brinson, M. and Brown, S. (eds) (1990) Ecosystems of the world. 15. Forested wetlands, Elsevier, Amsterdam, The Netherlands. McKee, K.L. (1993) Soil physicochemical patterns and mangrove species distribution – reciprocal effects? Journal of Ecology, 81, 477–487. McKee, K.L., Mendelssohn, I.A. and Hester, M.W. (1988) Reexamination of pore water sulphide concentrations and redox potentials near the aerial roots of Rhizophora mangle and Avicennia germinans. American Journal of Botany, 75, 1352–1359. Middelburg, J.J., Nieuwenhuize, J., Slim, F.J. and Ohowa, B. (1996) Sediment biogeochemistry in an East African mangrove forest (Gazi Bay Kenya). Biogeochemistry, 34, 133–155. Nedwell, D.B., Blackburn, T.H. and Wiebe, W.J. (1994) Dynamic nature of the turnover of organic carbon, nitrogen and sulphur in the sediments of a Jamaican mangrove forest. Marine Ecology Progress Series, 110, 223–231. Nickerson, N.H. and Thibodeau, F.R. (1985) Association between pore water sulfide concentrations and the distribution of mangroves. Biogeochemistry, 1, 183–192. Nielsen, T. and Andersen, F.Ø. (2003) Importance of leaf decomposition on phosphorus cycling in a mangrove sediment (submitted). Nielsen, O.I., Macintosh, D.J. and Kristensen, E. (2003a) Development of benthic metabolism and sulfur dynamics in deposited shrimp pond waste used as substratum for mangrove seedlings. (submitted to Aquaculture). Nielsen, O.I., Macintosh, D.J. and Kristensen, E. (2003b) Impact of fiddler crabs (Uca spp.) on rates and pathways of benthic mineralization in deposited shrimp pond waste. Submitted to Journal of Environmental and Marine Biology and Ecology, 289, 59–81. Odum, W.E. and Heald, E.J. (1972) Trophic analysis of an estuarine mangrove community. Bulletin of Marine Science, 22, 671–738. Ovalle, A.R.C., Rezende, C.E., Lacerda, L.D. and Silva, C.A.R. (1990) Factors affecting the hydrochemistry of a mangrove tidal creek, Sepetiba bay, Brazil. Estuarine Coastal and Shelf Science, 31, 639–650. Platong, J. (1998) Status of mangrove forests in Southern Thailand, New School Publishing, Pattani, Thailand.

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Poovachiranon, S. and Tantichodok, P. (1991) The role of sesarmid crabs in the mineralization of leaf litter of Rhizophora apiculata in a mangrove, southern Thailand. Phuket Marine Biology Centre Research Bulletin, 56, 63–74. Poovachiranon, S., Suraswadi, P., Kristensen, E. and Trakulsiripanich, C. (2003) Litter fall and litter transport in the Bangrong mangrove forest, Phuket, Thailand (in preparation). Primavera, H. (1993) A critical review of shrimp pond farming in the Philippines. Reviews in Fisheries Science, 1, 151–201. Revsbech, N.P., Madsen, B. and Jørgensen, B.B. (1986) Oxygen production and consumption in sediments determined at high spatial resolution by computer simulation of oxygen microelectrode data. Limnology and Oceanography, 31, 293–304. Rivera-Monroy, V.H., Day, J.W., Twilley, R.R., Vera-Herrera, F. and Coronado-Molina, C. (1995) Flux of nitrogen and sediment in a fringe mangrove forest in Terminos Lagoon, Mexico. Estuarine Coastal and Shelf Science, 40, 139–160. Rivera-Monroy, V.H., Madden, C.J., Day, J.W., Twilley, R.R., Vera-Herrera, F. and Alvarez-Guillen, H. (1998) Seasonal coupling of a tropical mangrove forest and an estuarine water column: enhancement of aquatic primary productivity. Hydrobiologia, 379, 41–53. Robertson, A.I. (1986) Leaf-burying crabs: their influence on energy flow and export from mixed mangrove forests (Rhizophora spp.) in north eastern Australia. Journal of Experimental Marine Biology and Ecology, 102, 237–248. Robertson, A.I. and Daniel, P.A. (1989) The influence of crabs on litter processing in high intertidal mangrove forests in tropical Australia. Oecologia, 78, 191–198. Robertson, A.I. and Phillips, M.J. (1995) Mangroves as filters of shrimp pond effluent: predictions and biogeochemical research needs. Hydrobiologia, 295, 311–321. Robertson, A.I., Alongi, D.M. and Boto, K.G. (1992) Food chains and carbon fluxes, in Tropical Mangrove Ecosystems (eds A.I. Robertson and D.M. Alongi), Coastal and Estuarine Studies 41, American Geophysical Union, Washington, DC, pp. 293–326. Sansanayuth, P., Phadungchep, A., Ngammontha, S., Ngdngam, S., Sukasem, P., Hishino, H. and Ttabucanon, M.S. (1996) Shrimp pond effluent: pollution problems and treatment by constructed wetlands. Water Science and Technology, 34, 93–98. Shunula, J.P. and Whittick, A. (1999) Aspects of litter production in mangrove from Unguja Island, Zanzibar, Tanzania. Estuarine Coastal and Shelf Science, 49, 51–54. Smith, T.J. III, Boto, K.G., Frusher, S.D. and Giddins, R.L. (1991) Keystone species and mangrove forest dynamics: the influence of burrowing by crabs on soil nutrient status and forest productivity. Estuarine Coastal and Shelf Science, 33, 419–432. Sundby, B., Gobeil, C.G., Silverberg, N. and Mucci, A. (1992) The phosphorus cycle in coastal marine sediments. Limnology and Oceanography, 37, 1129–1145. Suraswadi, P., Kristensen, E., Andersen, F.Ø., Holmer, M., Holmboe, N., Thongtham, N. and Flindt, M. (2003) Transport of nutrients in the Bangrong mangrove forest Phuket, Thailand; with emphasis on the impact of aquaculture (in preparation). Tam, N.F.Y. and Wong, Y.S. (1995) Mangrove soils as sinks for wastewater-borne nitrogen and phosphorus in mangrove soils. Hydrobiologia, 295, 231–241. Tam, N.F.Y. and Wong, Y.S. (1999) Mangrove soils in removing pollutants from municipal water of different salinities. Journal of Environmental Quality, 28, 556–564. Tanaka, K., Sugahara, K., Ohwaki, Y. and Choo, P.-S. (1998) C, N, P composition of suspended matter in Matang Mangrove Estuary, Malaysia. Japan Agricultural Research Quarterly, 32, 153–158. Thamdrup, B. (2000) Bacterial manganese and iron reduction in aquatic sediments, in Advances in Microbial Ecology, vol. 16 (ed. B. Schink), Kluwer Academic/Plenum Publishers, New York, pp. 41–84. Thamdrup, B. and Canfield, D.E. (2000) Benthic respiration in aquatic sediments, in Methods in Ecosystem Science (eds O.E. Sala, R.B. Jackson, H.A. Mooney and R.W. Howarth), Springer, New York, pp. 86–103.

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2

Coral reefs M.J. Atkinson and J.L. Falter

2.1

Introduction

The motivation for studying biogeochemical cycles of coral reefs has been to understand how these marine ecosystems maintain high carbon production, turnover and deposition of calcium carbonate in low-nutrient oceans (D’Elia & Wiebe, 1990). Over the past 80 years, a variety of ideas have been advanced to explain high carbon production. One of the prevalent views is that close physical relationships between heterotrophs and autotrophs (i.e. symbiosis) create an ecosystem where nutrients are either retained within the biota or recycled within the community. A more recent view, however, is that coral reefs produce mostly low quality organic carbon and inorganic carbon, requiring relatively little input of nutrients. The nutrients that are removed from the water column are exported as particulate material and recycled over spatial scales of at least hundreds of meters. All biogeochemical pathways, or reactions, described for pelagic systems occur in coral reef ecosystems (see D’Elia & Wiebe, 1990). The biogeochemistry of coral reef ecosystems, however, is different from pelagic systems in spatial arrangements, rates, stoichiometry and mechanisms that govern the reactions. These basic differences between reefs and pelagic systems are only recently becoming apparent. A major obstacle in advancing coral reef biogeochemistry has been that biogeochemical rates are normalized in different ways, and these rates are dependent on the scale of measurement. Biogeochemical rates in pelagic systems are normalized to biomass (g dry wt, Chl a, N or P content, ATP, etc.); measurements are extrapolated to larger scales by linear relationships with biomass per volume. This approach is widely practiced. In reef systems, however, it is problematic to measure biomass per area because biomass penetrates the carbonate substrate and is spatially very heterogeneous. Yet, rates are determined for biogeochemical processes and normalized to biomass making extrapolations to the whole system nearly impossible. Further, and more fundamentally, the quantity of biomass is not necessarily linearly related to biogeochemical rates. Many reactions are also normalized to surface area, but even this rather simple approach is unsatisfactory when applied. Different approaches and methods of normalization make it difficult to develop quantitative biogeochemical models. Many papers report organism-scale experiments in which transfers of biogeochemical compounds are described

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but not quantified; those papers are not referenced here because it is impossible to assess the quantitative significance of the processes they address. In a previous review, D’Elia and Wiebe (1990) described biogeochemical pathways and reported the organisms involved in each reaction; there were few rates reported, and those that were reported were per biomass. In this review, we try to advance the understanding of the basic processes controlling kinetics, and provide some estimates of rates per area; we do not want to repeat the information in that review. We briefly summarize the known biogeochemical reactions occurring in reefs and will then try to place them in the physical context of the reef with some estimates of the rates. We will also recommend, considering recent findings, a relatively new approach in making measurements of biogeochemical rates.

2.2

Coral reef morphology and zonation

Coral reefs are living structures that biogenically produce calcium carbonate; they maintain themselves at sea-level against the destructive forces of waves. Morphologies of reefs are shaped by their underlying geological structures, and the net growth, accretion and dissolution of carbonate structures (Grigg et al., 2002). It is necessary to define basic reef morphology (Stoddart, 1969) because biogeochemical reactions are related to different morphology and zones. The fore-reef is a transition region between the reef and the ocean, sloping from deeper shelf to the shallowest reef (Fig. 2.1). This slope can vary in steepness depending on the underlying topography and the upward growth of the reef. The fore-reef is usually comprised of spurs and grooves. Waves break and water moves through the grooves, constantly eroding the structure. The spur or buttress actively accretes, growing upward and outward to maintain the fore-reef (Fig. 2.1). The fore-reef rises to the reef crest which is the shallowest part of the reef. This area, classically, is comprised of coralline and turf algae which can be exposed at low tide. The reef flat is flat and shallow; but can vary in depth depending on whether the reef has been uplifted or sunk from island subsidence and from spatial variation in rates of reef accretion. Reef flats can be dominated by a variety of organisms, from pure stands of coralline algae to small coral, crustose coral, pure macro-algae and even pure soft corals and zooanthids. The community structure largely depends on wave energy (Grigg, 1998). Reef flats have been a convenient site for community scale metabolic studies; most research on biogeochemistry has been in this zone (Smith & Marsh, 1973). Back-reefs are comprised of sand deposits and coral knolls, often reticulated in wonderful patterns. The classic reef flat progrades to a lagoon which usually has patch reefs of varying sizes. These general morphological features are shown in Fig. 2.1. Fringing reefs on high

42

BIOGEOCHEMISTRY OF MARINE SYSTEMS

algal pavement reef crest

sand - rubble coral - algae

grooves

fore-reef slope

Fig. 2.1 Top: Aerial photograph of a section of the barrier reef of Chuuk Atoll, Marshall Islands (7°27′N, 151°51′E). Picture courtesy of the Coral Reef Research Foundation. Dark lines on the forereef indicate the presence of grooves. Bottom: Generalized sketch of fore-reef and reef flat for Bikini Atoll, Marshall Islands (11°30′N, 165°25′E) adapted from Munk and Sargent (1954).

islands show similar features with a fore-reef, reef flat, and sand deposition area; but no lagoon nor patch reefs. The basic morphological features and zones discussed above are readily apparent in air photos and even satellite images (Fig. 2.1). Generally, sand is bright (40–60% reflectance of incident light) making it easy to identify sand areas; algae and coral are darker, reflecting only 5–15% of light (Hochberg et al., 2002). Highly three-dimensional structures, including those dominated by coral, are very dark and surrounded by sand; note the dark regions identified

43

CORAL REEFS

in the air photo. Major algal taxa and coral are spectrally distinct, making it possible to map different bottom types (Hochberg et al., 2002). 2.3

Basic biogeochemistry

2.3.1

Carbon

In this section, we introduce the basic biogeochemical cycles of carbon, nitrogen, phosphorus, silica and iodine (Fig. 2.2). It is important to remember that NO2 4

DOP, DON, DOC

4

A

I, NH4, HPO4, Si(OH)4, NO3, CO2, HCO3

POMMixed

A

POMAllo

B

N2

C

D B 3

5

1

550CO2 + 30NH4 + HPO4 + xSi(OH)4

POMAuto + PQ*550O2 6

POMMixed OXIC

POMAll

6

NH4, HPO4, Si(OH)4, CO2

SUBOXIC

4NO3 + 5CH2O 2CH2O + SO4

8

ANOXIC 2CH2O

9

CaCO3

Ca + CO3

2

NO3

7

NH4, HPO4, Si(OH)4, CO2 4

2O2 + NH4

NO3 + H2O + 2H

5HCO3 + 2N2 + H + 2H2O

2HCO3 + HS + H CH4 + CO2

CaCO3

10

Ca + CO3

Transfer pathway

Reaction pathway

(A) (B) (C) (D)

(1) (2) (3) (4) (5) (6) (7) (8) (9) (10)

Uptake of dissolved nutrients Excretion – Respiration Particulate organic matter uptake Particulate organic matter export

Net primary production Nitrate assimilation N2 fixation Nitrification Calcification Organic matter remineralization Denitrification Sulfate reduction Methanogenesis Calcium carbonate dissolution

Fig. 2.2 Schematic of fundamental biogeochemical reactions occurring in coral reef systems. POM stands for particulate organic matter. Different types of POM are denoted by the subscripts Allo (allochthonous), Mixed (autochthonous), All (autochthonous + allochthonous), Auto (solely autotrophic).

44

BIOGEOCHEMISTRY OF MARINE SYSTEMS

the standing stocks of carbon, nitrogen and phosphorus in the water column are orders of magnitude lower than standing stock in the biota or in the sediments (Table 2.1). Table 2.1 is designed to show the mass per square meter for the water column a meter above the reef, the biota and the carbonate framework a meter below or into the reef. The meter scale is simply for ease in scaling and calculations. For example, one could increase or decrease the thickness of the water pool or the sediment pool to equal the mass in the biota. Or, if one wanted to estimate the flux (mmol m−2 d−1) of compounds through each pool, the pool size could be divided by a turnover time: i.e. turnover of water (104 per day for a current of 0.1 ms−1), turnover time of biota (10−2 per day for 100 day turnover) and accumulation of the sediment pool (by 15 × 10−6 per day for 5 mm per year accumulation). For example, phosphate flux in the water would be 5000 mmol m−2 d−1, phosphate flux in the biota would be 1 mmol m−2 d−1 Table 2.1 Mass of phosphorus (P), fixed nitrogen (N), and organic carbon (C) in mmol m−2 in 1 m of the water column above the benthos, in the living benthos (autotrophic and heterotrophic), and in the top-most 1 m of sediments (solid and dissolved phases) Pool 1 m water column Living benthos Autotrophic Heterotrophic 1 m sediments Solid phase Dissolved phase

P

N a

C

0.16–0.90

<7

a

10–250a

50b 50c

1350b ~2000i

22 400b ~5 000d

1400h(10 000)e,f,g,h <2j

34 000i <50 j

300 000h 5–125k

Phosphorus and fixed nitrogen include all inorganic and organic forms. Values for the water column are based on the data presented in Table 2.3. b Values for benthic autotrophs calculated from dry weight biomass estimates (Odum & Odum, 1955) and assuming a C:N:P ratio of 550:30:1 (Atkinson & Smith, 1983). c Benthic heterotroph-phosphorus content was calculated using dry biomass estimates (Odum & Odum, 1955) and assuming a phosphorus content of 1% by weight (Pomeroy & Kuenzler, 1969). d Benthic heterotrophic biomass content calculated assuming nearly all of the biomass is composed of CH2O. Organic phosphorus and total phosphorus (in parenthesis), and organic carbon calculated based on the data of e Entsch et al. (1983). f Atkinson (1987b). g Szmant and Forrester (1996). h Suzumura et al. (2002), assuming a porosity of 0.5 (Buddemeier & Oberdorfer, 1988) and a sediment density of 2.7 g cm−3. i Solid phase nitrogen content measured by Szmant and Forrester (1996). j Estimates for sediment dissolved pool sizes assume pore water dissolved inorganic phosphorus (DIP) concentrations <3 μM and dissolved inorganic nitrogen (DIN) concentrations <80 μM, a sediment porosity of 0.5 (see text for details) and that dissolved inorganic pools are much greater than dissolved organic pools. k This estimate assumes pore water dissolved organic carbon concentrations are equal to ambient water dissolved organic carbon concentrations (Tribble et al., 1990). a

45

CORAL REEFS

and in the sediments, 0.02 mmol m−2 d−1. One would conclude that phosphate flux through the water column is greatest and therefore is the largest potential source for most compounds. Carbon is removed from the water column by autotrophs and converted to both organic and inorganic carbon (Kinsey, 1985; Hatcher, 1997; Gattuso et al., 1998). Gross production is the total amount of carbon fixed over 24 h. It is typically estimated by adding day-time net photosynthesis (uptake of carbon and release of oxygen) to night-time respiration (release of carbon and uptake of oxygen), extrapolated over 24 h. This approach assumes light respiration equals dark respiration. Recent research at the Biosphere 2 coral reef (850 m2) has shown that light respiration can be twice as greater as dark respiration, making some of the overall numbers in the literature questionable (Langdon et al., 2002). In practice, oxygen is usually used to estimate net photosynthesis and respiration, and a photosynthetic and respiratory quotient of 1.0 is assumed (range 0.8–1.2; Kinsey, 1985). Many such measurements have produced a uniform metabolic standard for reefs (Kinsey, 1985). Gross community production varies from 100 to 2000 mmol C m−2 d−1, depending on habitat (Table 2.2). Low relief sand communities are the lowest (100–300 mmol C m−2 d−1) with reef flats being moderate (350–500 mmol C m−2 d−1) and high-relief communities of coral and algae showing maximal values (1000–2000 mmol C m−2 d−1). These overall metabolic rates of different reefs are consistent between reefs, suggesting that metabolic rates are independent of species composition. These basic communities can be identified easily from air photography and even satellite images, making ecosystem-wide estimates of gross production possible (Atkinson & Grigg, 1984; Andrefouet & Payri, 2000). Community respiration varies over the same range as gross production; the two are positively correlated (Table 2.2). Communities or bottom types with high production tend to exhibit high respiration, indicating much of the respiration of organic material occurs within the habitat, or more probably within the

Table 2.2 Gross primary production (P), Community respiration (R), Net Community Production (NCP), and Net Community Calcification (G) in mmol C m−2 day−1 for various communities as originally tabulated by Kinsey (1985) with additional data from Gattuso et al. (1993, 1996), Kraines et al. (1996, 1997), Boucher et al. (1998), and Andrefouet and Payri (2000). Means are in bold followed by the range in parentheses Habitat

P

R

NCP

G

Entire Reef-flat Algal pavement High coverage Sandy areas Shallow lagoon Entire Reef systems

640 (330–1580) 460 (170–580) 1180 (660–1920) 130 (80–230) 450 (210–1080) 390 (190–640)

600 (290–1250) 300 (40–560) 1280 (500–2000) 130 (90–200) 430 (180–790) 370 (190–570)

−220–310 0–130 −830–250 −40–30 −200–280 0–70

130 (20–250) 90 (70–110) 240 (110–320) 35 (10–70) 40 (20–55) 45 (3–135)

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

organism that produces the carbon (Falter et al., 2001). There are, however, patterns of net production and consumption on reefs. Net production, the excess carbon produced over 24 h period (gross production minus community respiration), varies spatially. In classic reef zonation (Odum & Odum, 1955; Kinsey, 1985), net carbon is produced on the fore-reef and algal crest, exported to the back-reef area as detritus and dissolved organic carbon (Fig. 2.1), sustaining and supporting a net heterotrophic community. In practice, the delineation is not so clear. Many back-reef areas that have significant rubble, can have high net production (Atkinson & Grigg, 1984). So much of this zonation is dependant on the nature of the substrate. Sand and muds have a tendency to be heterotrophic (gross production less community respiration is negative), while areas exposed to high water motion and hard substratum with algae tend be autotrophic (gross production less respiration is positive). Rich coral areas and knolls are usually close to a net production of zero (P/R of 1.0). The consensus has been that sustained inputs of nutrients from the ocean support net production, although past observations have repeatedly shown that nutrient concentrations change little across most fore-reefs and reef flats, and concentrations of particulate organic matter usually increase. The link between nutrient uptake and net carbon production has not been well established, and remains one of the more illusive aspects of coral reef biogeochemistry. Coral reef communities have also demonstrated the ability to take up suspended planktonic organic matter as a source of new carbon (Glynn, 1973; Johannes & Gerber, 1974; Ayukai, 1995; Sebens et al., 1997; Fabricius et al., 1998; Ribes et al., 1998; Yahel et al., 1998). Reported rates of particulate carbon uptake are low (<40 mmol C m−2 day−1) compared to rates of gross primary production and community respiration (Table 2.2), and have been demonstrated to be an unimportant source of carbon for some hard and soft coral communities (Ribes, 1998; submitted). However, the uptake of particulate organic matter may be an important source of food for some specific reef communities (Fabricius et al., 1998). There have been no studies measuring both carbon and nutrient fluxes in different morphological zones, nor studies relating net carbon sources and sinks to nutrient fluxes. For many reefs, errors in gross production and respiration are too high (10–15%) to get reliable estimates of net community production (Crossland et al., 1991) while changes in nutrient concentrations are nearly undetectable. Such experimental restrictions have limited our ability to understand the relationship between carbon and nutrient cycles in coral reefs. Carbon and nutrient budgets of lagoons and shallow water systems have been achieved by calculating water residence times with water or salt budgets (Smith, 1984). Atoll-wide budgets indicate that the C:P ratios of net production or consumption must be relatively high (near 300:1) for the C and P budgets to balance. Net carbon production on a large scale is close to zero (Crossland

CORAL REEFS

47

et al., 1991). The spatial arrangement of net carbon production and net carbon consumption in different communities, however, is not clear from these budgets. Sediments in coral reefs are typically less than 1% organic carbon, indicating little sink of carbon in these systems. A recent study by Falter et al. (2001) of Biosphere 2 coral reef mesocosm (completely closed nutrient cycles) revealed that measurements of production and respiration over periods shorter than one week do not adequately reflect longer term net ecosystem metabolism. The obvious suggestion is that most of the previous small scale studies, may in fact, be giving a very inadequate view of net carbon metabolism on reefs. In general, the spatial and temporal variation in net carbon metabolism of reefs is poorly understood. Rates of carbon deposition as calcium carbonate range from 10 to 20% of gross carbon production. Communities that have high organic production tend to have the highest calcification rates (Table 2.2). Calcification is positively correlated to light (Lough & Barnes, 2000). Recent research reveals that calcification is also positively correlated to saturation state of aragonite, the ion product of calcium and carbonate in water (Gattuso et al., 1999; Marubini & Atkinson, 1999; Langdon et al., 2000; Langdon, 2001). Decreases in carbonate ion concentration of tropical surface water in the next 60 years from rising atmospheric carbon dioxide may reduce calcification by upwards of 30% (Kleypas et al., 1999; Langdon, 2001). Dissolution of carbonates occurs inside coral heads (DiSalvo, 1971), interior pore-spaces of coral reef frameworks (Tribble et al., 1990) and from the erosion action of boring organisms. Rates of dissolution are believed to be much slower than rates of biogenic precipitation (<10%; Tribble et al., 1990). Thus carbon metabolism on reefs has a tri-modal distribution (Table 2.2), and estimates of carbon production and calcification can be made by basic knowledge of bottom type. The partitioning of net carbon production and consumption, and the relationships to nutrient cycles, however, are not well understood, nor characterized, especially under field conditions. Further advances in this area of research will delineate kinetic constraints on the biogeochemical rates and their links with the carbon cycle. 2.3.2

Dissolved organic matter

Dissolved organic carbon is ubiquitous in ambient waters over coral reefs and typically occurs at concentrations much greater than particulate organic matter (Table 2.1). Dissolved organic carbon and nitrogen is taken up and released by a variety of organisms (Schlichter & Liebezeit, 1991; Sorokin, 1991; FerrierPages et al., 1998; Ambariyanto & Hoegh-Guldberg, 1999; Hoegh-Guldberg & Williamson, 1999; Yahel et al., 2003). Rates of dissolved organic matter metabolism have been reported for specific organisms and can be either significant (Yahel et al., 2003) or insignificant to total metabolism (Ambariyanto &

48

BIOGEOCHEMISTRY OF MARINE SYSTEMS

Hoegh-Guldberg, 1999). Much of the metabolism of dissolved organic matter by benthic reef organisms has been attributed to symbiotic bacteria (FerrierPages et al., 1998; Yahel et al., 2003). Unfortunately, the cycling of dissolved organic matter at the community or ecosystem scale remains poorly understood because (1) rates of metabolism at these scales have not been measured or estimated, and (2) the chemistry of naturally occurring dissolved organic compounds in sea water is complex. Finally, the metabolism of specific organic compounds occurring in low concentrations, such as steroids, can have substantial impacts on organismal biology (Tarrant, 2002). 2.3.3

Nitrogen

Dissolved species of nitrogen in the water column include all oxidation states of nitrogen: nitrate, nitrite, ammonium and organic nitrogen compounds (Table 2.3). Typical concentrations of nitrate and ammonium vary by an order of magnitude, from 0.05 to 0.5 μM N, while nitrite is usually below 0.1 μM N. Dissolved organic nitrogen is the largest pool in the water column usually near 5 μM but can be 10-fold higher. Particulate organic nitrogen concentrations are also larger than the inorganic compounds, near 1 μM N. The amount of nitrogen in biomass and sediments, however, is two to three orders of magnitude higher than the nitrogen in the water column (Table 2.1). Table 2.3 Concentrations of dissolved and particulate nutrients in μM (mmol m−3) reported from waters around coral reefs, presented as both the range of reported values and range of typical values (as interpreted by the authors) Nutrient

Range

Typical

Reference

DIP NO3− NO2− NH4+ Si(OH)4 DOP DON DOC POP PON POC

0.01–0.6 0.05–9.8 ud–0.3 ud–2.4 0.7–11.7 0.03–1.6 ud–22.0 9–290 0.06–0.3 0.1–2.5 2–127

0.05–0.3 0.05–0.5 <0.1 0.05–0.5 <5 0.05–0.3 <5 ~100 0.06–0.3 <1.0 ~10

a, b, c, e, g, h, i, j, k, l, m, q b, c, d, e, g, h, i, j, m, n, q d, h b, c, d, e, g, h, i, j, m, n, q b, h, i, m a, e, h, i b, e, g, h, i, j, n, p f, g, j, o, p e b, e, g, n a, b, f, g, o

DOP, DON, and DOC represent dissolved organic phosphorus, nitrogen, and carbon, respectively. POP, PON, and POC represent particulate organic phosphorus, nitrogen, and carbon, respectively. “ud” = undetectable. (a) Atkinson, 1987b; (b) Crossland and Barnes, 1983; (c) Crossland et al., 1984; (d) Falter and Sansone, 2000; (e) Furnas et al., 1995; (f) Gordon, 1971; (g) Johannes et al., 1972; (h) Johannes et al., 1983; (i) Kinzie III et al., 2001; (j) Lapointe, 1997; (k) Pilson and Betzer, 1973; (l) Sargent and Austin, 1949; (m) Tribble et al., 1990; (n) Webb et al., 1975; (o) Yahel et al., 2003; (p) Alongi et al., 1996; (q) Szmant and Forrester, 1996.

49

CORAL REEFS

The redox reactions of nitrogen that occur on reefs are similar to those that occur in pelagic systems: including nitrogen uptake and assimilation, reduction, nitrification and nitrogen fixation (see Fig. 2.2). Reef autotrophs take up ammonium, nitrate and nitrite from the water column, reduce them to organic nitrogen compounds and assimilate them into biomass. Ammonium and nitrate removal from the water, or uptake, is limited or controlled by diffusion through diffusive boundary layers between the benthic autotrophs and the bulk water. Rates of ammonium uptake (mmol N m−2 d−1) are proportional to concentration (Thomas & Atkinson, 1997). The proportionality constant S (m d−1) is positively correlated to water velocity and the roughness of the community, and appears to be directly related to the dissipation of energy in the water (see later section on hydrodynamic control of biogeochemical cycles). Values of S have only recently been reported in the literature; they range 2–16 m d−1 and are summarized in Table 2.4. Uptake of nitrite has not yet been studied.

Table 2.4 Uptake (A) and Release (B) rates of DIP, ammonium, and nitrate in mmol m−2 day−1; and (C) uptake rate coefficients (S, where rate = S × Concentration) for each nutrient in m day−1 DIP

NH4+

NO3−

Reference

– – – 7.8 ± 4.5 0.1–4.8 0.8–2.3

– – – – 0.2–5 0.8–2.3

Pilson and Betzer (1973) Atkinson (1987c) Atkinson and Bilger (1992) Steven and Atkinson (2003) Atkinson et al. (2001) Falter (2002)

– ns 4

Steven and Atkinson (2003) Atkinson et al. (2001) Tribble et al. (1990)*

– – – – – – – 4.4 ± 0.4 15.5 ± 2.1

Pilson and Betzer (1973)* Atkinson (1987c) Atkinson and Bilger (1992) Bilger and Atkinson (1995) Thomas and Atkinson (1997) Larned and Atkinson (1997) Steven and Atkinson (2003) Atkinson et al. (2001) Falter (2002)

(A) Uptake 0.9 ± 0.2 0.2–1.1 1.1 ± 0.2 1.2 ± 0.6 0–0.8 0.7–1.0 (B) Release 0.9 0.5 ± 0.3 0.25

4.3 2.1 ± 1.4 2

(C) S 4.5 9 1.2–15 0.4–5.2 – 2.9–4.1 5.8 ± 3.4 7.4 ± 2.0 9.3 ± 1.3

– – – 2.2–12.7 1.4–12.4 6.1–12.7 11.1 ± 6.4 7.5 ± 2.1 15.5 ± 2.1

Values are single values, ranges, or means with errors, depending on the source of data. See details of publications for determination of errors. Publications reporting significant uptake or release, but no rates are not listed in this table; citations to those publications are in cited literature. (*) represents data that has been reinterpreted. ns = not significant.

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

When ammonium is elevated above ambient, nitrate increases in the water, demonstrating a large potential for nitrification. We have routinely observed that 10% of ammonium uptake is oxidized to nitrate within hours of uptake (Atkinson et al., unpublished data). There is very little work showing rates of nitrification. Denitrification rates in reef sediments are estimated to be 1–4 mmol N m−2 d−1 (see discussion of interstitial geochemistry below). Given rate constants of about 15 m d−1 for ammonium and a concentration of 0.4 mmol N m−3 for ammonium and nitrate, give an uptake of nitrogen of only about 6 mmol N m−2 d−1. This rate of nitrogen uptake is low considering the amount of phosphate removed and the CNP ratio of biomass (see later discussion). The remaining nitrogen required to balance these budgets could be from particles or from nitrogen fixation. There is some indication that particles can support a large fraction of nitrogen (up to 50%) taken up by the reef (Erez, 1990; Ribes et al., 2003), however, these fluxes are not well established. The extent to which nitrogen fixation supplies this missing nitrogen is still undecided. Areal rates of nitrogen fixation on reefs vary by 2–3 orders of magnitude (Smith, 1984; Wilkinson et al., 1984; Atkinson, 1988; Williams & Carpenter, 1998). Larger scale experiments are required to address community scale nitrogen fixation. Williams and Carpenter (1998) have shown that rates of acetylene reduction (a proxy for N2-fixation) increase with flow velocity, suggesting that many of the experiments in the literature may not be properly characterizing the magnitude of the flux. Typically, dissolved organic nitrogen is exported from reef communities (Wiebe et al., 1975; Wilkinson et al., 1984); perhaps as high as ~10% of the total uptake of nitrate and ammonium. The nature of this dissolved organic nitrogen and the rate kinetics of uptake are also not known, so it is difficult to attach rates of uptake or recycling to this group of compounds. We believe, however, that the dissolved organic compounds may be fairly refractory and require a bacterial or sponge community for remineralization, so uptake will become largely dependent on the community structure of the benthos. 2.3.4

Phosphorus

Phosphorus occurs entirely as either inorganic or organic phosphate. Typical concentrations of dissolved inorganic phosphate in the water column vary from near undetectable concentrations to 0.3 μM. Typical concentrations of dissolved organic phosphate and particulate organic phosphate have a similar range to that of dissolved inorganic phosphate (Table 2.3). Alkaline phosphatase, an extra-cellular enzyme that cleaves the phosphate mono-ester bond in organic molecules, is probably ubiquitous in reef communities (Atkinson, 1987a). There are, however, no studies that calculate an actual rate of phosphate flux from dissolved organic phosphate to the reef surface; making it impossible to determine whether there is any significant turnover through the dissolved

CORAL REEFS

51

organic phosphate pool. It is likely that typical rate constants for mono-ester bonded phosphate will be similar to rate constants for the inorganic pool. It is widely known that phosphate adsorbs to calcium carbonate at high concentrations (Simkiss, 1964), however, there is no evidence that this is a significant pathway in coral reefs. Uptake of phosphate, even onto sand, is controlled by organisms and is highly correlated to respiration rates (Atkinson, 1987c). Sand exposed to mercuric chloride or formaldehyde releases large quantities of phosphate, demonstrating that large amounts of phosphate are stored in the biota. Phosphate that is assimilated into calcifying organisms gets incorporated into skeletal material at concentrations about 5–10 μmol P g−1 CaCO3. Rate constants or S values, range 1–9 m d−1 (Table 2.4), but in the field, uptake of phosphate is close to being mass transfer limited (see discussion below), as is ammonium. For typical field values of S (9 m day−1), rates of phosphate uptake from the dissolved pool are ~1 mmol P m−2 d−1 (Table 2.4). A substantial percentage (up to 30%) of this material is incorporated into calcium carbonate and is a sink in the sediments, creating low phosphate environments in lagoons with restricted water circulation (Smith, 1984). The remaining phosphate is exported as particulate or dissolved organic phosphate to adjacent communities. It is apparent that both the C:P ratio of organic material and the C:P ratio in carbonates can change in shallow-water, low-phosphate environments (Atkinson, 1987b), but it is not clear whether this occurs within different habitats of coral reefs, from differential uptake rates based on hydrodynamics. 2.3.5

Silica

There are no published studies of silica on coral reefs. Silica minerals are not abundant on coral reefs, so the biogeochemistry of silica has been ignored. It is evident, however, from the results of our uptake experiments (we measure silica in all our experiments) that silica concentrations in the water sometimes decreases and sometimes increases. This result is probably related to whether benthic diatoms are occupying the reef surfaces. At Biosphere 2 coral reef mesocosm, silica goes through a distinct seasonal cycle related to blooms of benthic diatoms (Atkinson et al., 2001). Concentrations of silica in pore-water can also change seasonally (Falter, 1998). 2.3.6

Iodine

Iodine is a basic nutrient compound in the ocean. Iodate (IO3) is the oxidized form and abundant in ocean. Iodate is removed by plants, iodide I− is released and oxidizes chemically and bacterially to iodate. The biogeochemical cycle has not been studied in reef ecosystems, but we believe it will be one of the most interesting. It appears iodate can be removed from the water column by reef communities very rapidly and has rate constants similar to nitrate (Atkinson,

52

BIOGEOCHEMISTRY OF MARINE SYSTEMS

unpublished data). Marine aquarists have also discovered that daily to weekly additions of potassium iodide promote coral growth, especially in aquaria with high nitrate (Atkinson, in preparation). The relationships between iodide uptake, thyroid hormones and calcification have not yet been studied, but we suspect it will be a major development in determining the health of coral reefs.

2.4

Interstitial geochemistry and hydrology of coral reef frameworks

Coral reef frameworks are partially lithified carbonate structures on which reef communities grow. Pore-waters of reef frameworks are mostly anaerobic and contain elevated levels of dissolved nutrients (Skyring & Chambers, 1976; Buddemeier & Oberdorfer, 1983; Corredor & Capone, 1985; Sansone, 1985; Sansone et al., 1990). This is a common feature of many coral reefs (Sansone et al., 1990), and a direct result of oxidation of organic matter (Sansone et al., 1990; Tribble et al., 1990). The subsequent production of carbonic acid from the oxidation of organic matter lowers pore-water pH and reduces the activity of the carbonate ion, thus facilitating in situ dissolution of carbonate (primarily aragonite) (Sansone et al., 1990; Tribble et al., 1990). Reef pore-waters become anoxic at depths less than a meter into the framework, however, some reef pore-waters become anoxic within centimeters of the framework surface (Entsch et al., 1983; Haberstroh & Sansone, 1999; Falter & Sansone, 2000). Pore-water nutrient concentrations typically vary up to 3 μM DIP, 6 μM NO2− + NO3−, 80 μM NH4+, 100 μM Si(OH)4 and 3.5 mM dissolved inorganic carbon (DIC) (Buddemeier & Oberdorfer, 1983; Entsch et al., 1983; Corredor & Capone, 1985; Sansone, 1985; Sansone et al., 1990; Tribble et al., 1990; Szmant & Forrester, 1996; Haberstroh & Sansone, 1999; Falter & Sansone, 2000; Suzumura et al., 2002), although much higher concentrations can be found in more diagenetically altered pore-water (e.g. Florida Keys: Szmant & Forrester, 1996; Sansone et al., unpublished data). The hydraulically driven transport of water into, through, and out-of coral reef frameworks has long been hypothesized as the primary mechanism providing particulate organic matter to drive framework diagenesis. Rates of organic carbon oxidation indicate that a continual supply of organic matter is necessary to support the metabolism of the framework (Tribble et al., 1990). Hydraulically driven advection of water can force organic particles into permeable sediments (Huettel et al., 1996; Huettel & Rusch, 2000). Unfortunately, the exact origins, transport, and fate of reef framework organic materials have not yet been determined. Coral reef frameworks are very permeable with typical hydraulic conductivities (K) ranging between 10 and 1000 m day−1 (Buddemeier & Oberdorfer, 1988; Roberts et al., 1988; Tribble et al., 1992; Falter & Sansone, 2000). Consequently, ambient pressure gradients across reef frameworks are large enough to drive the flow of interstitial water.

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Wave-induced mixing is considered an important process driving the exchange of shallow (<1 m within framework) pore-water with overlying water (Haberstroh & Sansone, 1999; Falter & Sansone, 2000). Haberstroh and Sansone (1999) showed that cross-reef variations in the vertical redox structure of pore-waters within the top two meters of the Checker Reef framework, Oahu, Hawaii were affected by the direction of waves impinging on the reef. Falter and Sansone (2000) showed that dissolved oxygen within the top half meter of the Checker Reef framework increased with both average wave height and average permeability of the sediments. Very low concentrations of nitrous oxide and nitrite (reaction intermediates for nitrification and denitrification) in shallow pore-waters of Checker Reef indicate that suboxic processes are limited to small spatial scales or microzones (Falter & Sansone, 2000). Elevated methane concentrations in anoxic reef pore-waters with un-depleted sulfate concentrations indicate that microzones are present in the deeper and more anoxic zones of reef frameworks as well (Sansone, 1985; Sansone et al., 1990). The residence time of pore water, 1 m deep in the framework of Checker Reef, Oahu, Hawaii is estimated to be ~2 days, while the residence time of pore-waters 2 m deep in the same framework is ~40 days (Tribble et al., 1992). This difference in residence time is consistent with the idea that any waveinduced pore-water exchange will naturally attenuate with depth in the framework (Haberstroh, 1994). Using residence time of pore-waters at 1 m, pore-water concentrations reported by Tribble et al. (1990), an effective porosity of 0.25 (Buddemeier & Oberdorfer, 1988), yields estimates of net framework nutrient fluxes for Checker Reef of 0.25 mmol P m−2 day−1 and 3.7 mmol N m−2 day−1. The ratio of N:P in shallower, oxic pore waters is roughly one-half to one-third the ratio in deeper pore-waters (Haberstroh & Sansone, 1999; Falter & Sansone, 2000). This change in N:P ratio may indicate the importance of denitrification in decreasing the potential flux of dissolved inorganic nitrogen out of the Checker Reef framework by a few mmol N m−2 day−1. More direct estimates of denitrification in other coral reefs have provided comparable rates of 1–4mmol N m−2 day−1 (Corredor & Capone, 1985; Seitzenger & D’Elia, 1985). Hydraulic conductivities reported for Checker Reef are similar to the range of conductivities reported for other reefs (K = 15–250 m day−1; Tribble et al., 1992; Falter & Sansone, 2000). However, Checker Reef lacks the hydraulic resistance of a lithified surface reef plate with relatively low conductivity (K ≤ 2 m day−1; Buddemeier & Oberdorfer, 1988). Such reef plates are commonly found capping many reef frameworks, restricting the exchange of pore-waters within the framework with the overlying sea-water. Checker Reef is also somewhat protected from direct exposure to oceanic swell by a large barrier reef; so the hydraulic forces driving pore-water exchange and transport within this framework may be weaker than average. Nevertheless, fluxes of biogeochemical compounds from interstitial pore-waters appear to be related to

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hydrodynamics and a significant proportion of fluxes from the water column to the reef benthos. We believe results from Checker Reef are representative of many back reef environments which are characterized by low energy wave conditions and high sand and rubble coverage.

2.5

Mass transfer-limited biogeochemical rates

Macrophytes, corals, and reef surfaces with turf algae have high potentials for assimilation of nutrients. Typically, the rates of uptake are maximal, or saturated, at water concentrations 100–1000 fold higher than ambient concentrations (Atkinson, 1988) indicating a very great demand for nutrients. At relatively low nutrient concentrations, diffusion of the nutrient compounds through water becomes the rate limiting step for uptake (Bilger & Atkinson, 1992); and a nutrient-depleted boundary layer, called the diffusive boundary layer, forms adjacent to all surfaces (or elevated in the case for nutrient release). The rate of uptake is controlled by diffusion of the nutrient molecules across the diffusive boundary layer. This rate is termed the mass transfer-limited rate and is proportional to the gradient in nutrient concentration between the water and the reef surface, and the thickness of the concentration boundary layer. If the demand for nutrients at these surfaces is very high, which is typical, then the concentration at the surface is negligible compared to the concentration in the water and the mass transfer-limited rate is proportional to the bulk concentration (rate = S × concentration; where S is rate constant with units m d−1; when concentration is mmol m−3, the rate is mmol m−2 d−1). Rates of uptake and release can be estimated using mass transfer correlations from the engineering literature (Bilger & Atkinson, 1992; Thomas & Atkinson, 1997). We report references of measured and calculated nutrient uptake and release rate constants, S (m d−1), for reef communities under different amounts of nutrient loading and water velocities (Table 2.2). Nutrient uptake by natural assemblages of corals and algae appears to be mass transfer limited (Baird & Atkinson, 1997; Falter, 2002; Steven & Atkinson, 2003). A mass transfer equation for nutrient uptake or release is Rate = (St × Ub) (Cb − Cw), where St is the Stanton number, Ub is water velocity and Cb and Cw are the concentrations of nutrients in the bulk water outside the boundary layer (Cb), and at the wall or surface (Cw), which is assumed to be negligible. The Stanton number (S/Ub) can be thought of as a physical descriptor of the bottom, characterizing the maximum potential rate of uptake at a given flow speed. It is also a dimensionless ratio of the flux into, or out-of the reef, divided by flux past the bottom for a meter water column. Stanton numbers for reef communities are of order 10−4–10−3. Thus the amount of nutrient removed from flowing water is a very small percentage of the total amount of nutrients flowing past the reef benthos (Atkinson & Bilger, 1992). For example, for water 1 m deep flowing at

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0.2 m s−1 across a 100 m transect with a phosphate concentration of 0.1 μM and a rate constant of 9 m day−1, the uptake rate is 9 m d−1 × 0.1 μM or ~1 mmol P m−2 d−1. Given the 8-minute time-period required for the water to flow across this transect, only 5% of the phosphate in the water column would be removed. It is important to remember that rate constants of uptake and release of biogeochemical compounds are limited by physical processes and confined to relatively low rates. Thus, any significant recycling of compounds – uptake and release – must occur on scales of at least hundreds of meters and more like kilometers. Even compounds that are ejected into the water column by animals advect and diffuse, mixing much faster than they are removed from the water column (Atkinson & Smith, 1987). Values of S decrease when communities of reef autotrophs are loaded with nutrients for several days at rates approaching about 10 times the normal rates in the field (Bilger & Atkinson, 1995). It is not clear whether the uptake kinetics of the plants actually slow down increasing the surface concentration (Cw), or whether excretion increases (Bilger & Atkinson, 1995). The mass transfer relationships from the engineering literature have been a convenient way of parameterizing uptake and release, however, we believe the fundamental control is the dissipation of energy over the reef (Hearn et al., 2001), which can be measured directly. The bottom exerts a force opposing the flow, called shear stress. A reorganization of the mass transfer equations from the engineering literature shows the rate constant S for nutrient uptake is directly linearly proportional to the shear stress to the 0.4 root (Hearn et al., 2001; Falter, 2002), demonstrating the simplicity of this approach (Fig. 2.3). The friction of the bottom creates turbulent eddies above the reef and the small-scale eddies can be used to calculate the dissipation of energy by the bottom. We do not know to what extent turbulence from external sources can drive mass transfer – this is now a subject of ongoing research. Recent work has shown that oscillating flows from waves actually create higher mass transfer for a given velocity than steady flows (increase 1.5–2.0 fold; Falter, 2002). Thus, reefs and reef communities that have more shear stress, or experience higher levels of dissipation through turbulent eddies, we believe, will transfer biogeochemical compounds faster through the water column. It is apparent that nutrient uptake into a coral reef flat is mass transfer limited (Falter, 2002); these mass transfer limited rates are maximal rates. Interestingly, the uptake of both nitrogen and phosphorus compounds is still relatively low compared to the production of organic carbon, giving N:P ratios of net production of around 10:1 (Atkinson & Bilger, 1992; Falter, 2002). C:N:P ratios of plants on reefs are also relatively high compared to the Redfield ratio and near something like 550:30:1 (Atkinson & Smith, 1983). Nitrogen is missing. At Biosphere 2 mesocosm, a completely enclosed re-circulating system, the nitrogen content in the water has increased and phosphate decreased such that the uptake based on mass transfer limited rates

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20 r2 = 0.94

–1

S (m day )

15

10

gypsum corals high relief rubble

5

low relief rubble P. compressa P. damicornis

0 0

1

2 0.4

τw0.4 (N

m

3

– 0.8

)

Fig. 2.3 Mass transfer coefficient for various communities involving either the uptake of ammonium (Thomas & Atkinson, 1997) or release of calcium from the dissolution of gypsum-coated corals (Baird & Atkinson, 1997) versus benthic shear stress raised to the 0.4 power.

matches the C:N:P in biomass (Atkinson et al., 2001). Real reefs in which the dissolved nutrients of nitrogen and phosphorus are similar in concentration (Table 2.3) cannot escape this tendency to be nitrogen deprived. Clearly, input from particles and nitrogen fixation must be important to the overall nitrogen budget. Areas of reefs with higher dissipation of energy, namely fore-reef and reef flats may have higher rates of nutrient input, lower C:N:P ratios and better quality organic carbon for food. We suggest that the spatial arrangement of net autotrophy and heterotrophy must be on scales of hundreds of meters and have characteristic differences in water flow and dissipation of energy (Falter, 2002). These ideas are yet to be proven. 2.6

Coral growth in high nutrient water

Nutrient loading and its subsequent impact is one of the more important issues concerning conservation and protection of coral reefs. It is widely

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believed that any nutrient input to coral reefs is deleterious. The argument is actually based on an incorrect historical view of how coral reefs recycle nutrients. Nutrient concentrations were observed to change very little across the relatively narrow reef flats of the Indo-Pacific (100–300 m). Scientists chose to believe that nutrients were recycled rapidly through the water column on scales of several meters (see Hearn et al., 2001). The notion developed that reefs recycle their nutrients through biologically mediated mechanisms, and any input of nutrients altered or perturbed these processes. Corroborating this view were early results that nutrients, in particular phosphate, affected community metabolism (Kinsey & Davies, 1979); and elevated nitrate and ammonia retard coral growth. More recently, nutrient loading appears to alter the reproductive patterns in corals (Ward & Harrison, 2000; Koop et al., 2001). Furthermore, nutrients actually stimulate the specific growth rate of many macro-algae; so sustained nutrients add to algal growth with the possible overgrowth of coral. Thus, nutrients are presently recognized as deleterious. The above conclusion, that nutrients are deleterious to a reef ecosystem, is simply incorrect. First, it is now known, based on the rate constants for mass transfer limitation of nutrient uptake, that nutrients cannot be recycled through the water column on such small scales of a reef flat; it is simply physically impossible. Thus, it is very difficult to perturb any biogeochemical cycles or biologically mediated mechanism with relatively small additions of nutrients to the water column. The reef barely can take up these nutrients. Reef communities must also take up very large quantities of nutrients to change the amount of nutrients in the biomass (Table 2.1). For example, most non-significant responses in the enrichment on coral reefs (ENCORE) experiments were indeed from too small a nutrient loading (Koop et al., 2001). Corals must be exposed to very high nutrient concentrations (at least 10 fold above ambient) for many days before there is a negative response in growth (Dubinsky & Jokiel, 1994; Marubini & Atkinson, 1999). Corals also can, and do, grow well at elevated nutrient concentrations, but it is not understood how other biogeochemical compounds influence these rates (Atkinson et al., 1995). Maintaining high concentrations of nitrate or ammonium in water over coral reefs is difficult. Usually, relatively rapid horizontal mixing from currents is much greater that the relatively slow uptake to the bottom (Table 2.4). The only significant anthropogenic impacts are in areas of groundwater or surface water discharge onto shallow reef flats where relatively large areas are impacted by nutrient input (Naim, 1993). In the presence of herbivory, it also has been difficult to show that algae even over-grow corals (Jompa & McCook, 2002), or are deleterious to their growth (McCook, 1999; McCook et al., 2001). We need more studies in which the organism or community responses are a function of actual nutrient loading, per area of benthos, not a function of concentration. There are far

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too many experiments in the literature that just report the concentrations of nutrients in an experiment instead of the flux of nutrients into the benthic organisms, or the increase of nutrients in internal pools. At present, any careful research on the effects of nutrients on coral growth must consider the nature of experimental apparatus and its flow characteristics, carbonate chemistry and light. Impacts to coral reefs from nutrients are probably indirect and long-term on an ecosystem scale (Smith & Buddemeier, 1992). Higher nutrient loading probably stimulates net production, creating ever increasing pools of organic matter and detritus. If not removed from the system by waves and currents, it is plausible that nutrients would encourage a bacterial fauna that promotes disease vectors, increasing the susceptibility and spread of disease. Coral diseases must be studied under reasonable nutrient fluxes. Some reefs experience high nutrients, high net production and high transport of organic matter to the shelf (notably fringing reefs of Japan and Western Australia), whereas other reef tracts receiving large amounts of nutrients collect organic matter (Florida Keys). Thus, the effects of nutrients must be interpreted on a local scale with respect to local biogeochemical and hydrodynamic budgets, and making sweeping conclusions that all nutrient input is deleterious, is simply irresponsible.

2.7

Measurement techniques

Many experiments include measurements of biogeochemical rates in small incubation chambers, reporting concentrations and a response by the experimental organism. Water is mixed by some arbitrary stirring mechanism. The underlying rationale for this approach is the belief that concentration is the only parameter controlling nutrient uptake (Atkinson, 1988). While metabolic responses can be described using hyperbolic relationships as a function of nutrient concentration (Michaelis-Menton type kinetics, Monod, etc.), these mathematical descriptions do not capture the underlying physics that controls nutrient uptake. The results from these types of experiments are subject to variability in experimental design and method of normalization of the rates. Most rates determined from these small-scale jar-type experiments are sufficient for getting nutrients into the autotrophs, but are simply not correct to apply to the field. We believe the correct parameterizations for most of these processes are not yet discovered. We still do not have appropriate facilities for maintaining and studying biogeochemistry of coral reefs. The understanding of their biogeochemistry will not advance unless appropriate flumes and mesocosms are constructed. It will take some national or international shared research facility at the appropriate place to make such advances.

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Summary statements

(1) All biogeochemical reactions found in pelagic ecosystems are also found in coral reef ecosystems. (2) The stoichiometry and mass balance relationships between organic and inorganic carbon production, and nutrient cycles are not well known. (3) Nutrient uptake and release rates in reefs occur at near mass transfer limits and may show water velocity dependence, depending on the reef, and its hydrodynamic characteristics. (4) Organic carbon inputs and remineralization within sediment interstitial spaces are related to hydrodynamic forces, and may be quite variable depending on variation in wave fields, large-scale pressure gradients, and the hydraulic permeability of the framework. (5) Generally, physical limits of nutrient inputs create relatively high C:N:P ratios of organic material. High C:N:P ratios and high dissipation of energy explains high carbon production in these systems. (6) Recycling of dissolved compounds through the water column must occur on scales of at least hundreds of meters and not on smaller scales. (7) Experiments must consider community size and the nature of the experimental environment to achieve reasonable estimates of biogeochemical fluxes. References Alongi, D.M., Tirendi, F. and Guldrick, A. (1996) Organic matter oxidation and sediment chemistry in mixed terrigenous-carbonate sands of Nigaloo Reef, Western Australia. Marine Chemistry, 54, 203–219. Ambariyanto and Hoegh-Guldberg, O. (1999) Net uptake of dissolved free amino acids by the giant clam, Tridacna maxima: alternative sources of energy and nitrogen? Coral Reefs, 18, 91–96. Andrefouet, S. and Payri, C. (2000) Scaling-up carbon and carbonate metabolism of coral reefs using in-situ data and remote sensing. Coral Reefs, 19, 259–269. Atkinson, M.J. (1987a) Alkaline phosphatase activity of coral reef benthos. Coral Reefs, 6, 59–62. Atkinson, M.J. (1987b) Low phosphorus sediments in a hypersaline marine bay. Estuarine Coastal and Shelf Science, 24, 335–348. Atkinson, M.J. (1987c) Rates of phosphate uptake by coral reef flat communities. Limnology and Oceanography, 32, 426–435. Atkinson, M.J. (1988) Are coral reefs nutrient-limited? 6th International Coral Reef Symposium, Townsville, 1, 157–166. Atkinson, M.J. and Bilger, R.W. (1992) Effect of water velocity on phosphate uptake in coral reef-flat communities. Limnology and Oceanography, 37, 273–279. Atkinson, M.J. and Grigg, R.W. (1984) Model of a coral reef ecosystem: II. Gross and net benthic primary production at French Frigate Shoals, Hawaii. Coral Reefs, 3, 13–22. Atkinson, M.J. and Smith, D.F. (1987) Slow uptake of 32P over a barrier reef flat. Limnology and Oceanography, 32, 436–441. Atkinson, M.J. and Smith, S.V. (1983) C:N:P ratios of benthic marine plants. Limnology and Oceanography, 28, 568–574.

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Atkinson, M.J., Carlson, B. and Crow, G.L. (1995) Coral growth in high-nutrient, low-pH seawater: a case study of corals cultured at the Waikiki Aquarium, Honolulu, Hawaii. Coral Reefs, 14, 215–223. Atkinson, M.J., Falter, J.L. and Hearn, C.J. (2001) Nutrient dynamics in the Biosphere 2 coral reef mesocosm: water velocity controls NH4 and PO4 uptake. Coral Reefs, 20, 341–346. Ayukai, T. (1995) Retention of phytoplankton and planktonic microbes on coral reefs within the Great Barrier Reef, Australia. Coral Reefs, 14, 141–147. Baird, M.E. and Atkinson, M.J. (1997) Measurement and prediction of mass transfer to experimental coral reef communities. Limnology and Oceanography, 42, 1685–1693. Bilger, R.W. and Atkinson, M.J. (1992) Anomalous mass transfer of phosphate on coral reef flats. Limnology and Oceanography, 37, 261–272. Bilger, R.W. and Atkinson, M.J. (1995) Effects of nutrient loading on mass-transfer rates to a coral-reef community. Limnology and Oceanography, 40, 279–289. Boucher, G., Clavier, J., Hily, C. and Gattuso, J.P. (1998) Contributions of soft-bottoms to the community metabolism (primary production and calcification) of a barrier reef flat (Moorea, French Polynesia). Journal of Experimental Marine Biology and Ecology, 225, 269–283. Buddemeier, R.W. and Oberdorfer, J.A. (1983) Pore-water biogeochemistry in a coral reef. EOS, 64, 735. Buddemeier, R.W. and Oberdorfer, J.A. (1988) Hydrogeology and hydrodynamics of coral reef pore waters. Sixth International Coral Reef Symposium, Australia, 2, 485–490. Corredor, J.E. and Capone, D.G. (1985) Studies on nitrogen diagenesis in coral reef sands. Fifth International Coral Reef Congress, Tahiti, 3, 395–399. Crossland, C.J. and Barnes, D.J. (1983) Dissolved nutrients and organic particulates in water flowing over coal reefs at Lizard Island. Australian Journal of Marine and Freshwater Research, 34, 835–844. Crossland, C.J., Hatcher, B.G. and Smith, S.V. (1991) Role of coral reefs in global ocean production. Coral reefs, 10, 55–64. Crossland, C.J., Hatcher, B.G., Atkinson, M.J. and Smith, S.V. (1984) Dissolved nutrients of a highlatitude reef, Houtman Abrolhos Islands, Western Australia. Marine Ecology Progress Series, 14, 159–163. D’Elia, C. and Wiebe, W. (1990) Biogeochemical nutrient cycles in coral reef ecosystems, in Coral Reefs: Ecosystems of the World Series (ed. Z. Dubinsky), Elsevier Science Publishers, Amsterdam, pp. 49–74. DiSalvo, L.H. (1971) Regenerative function and microbial ecology of coral reefs. II. Oxygen metabolism in the regenerative system. Canadian Journal of Microbiology, 17, 1091–1100. Dubinsky, Z. and Jokiel, P.L. (1994) Ratio of energy and nutrient fluxes regulates symbiosis between zooxanthellae and corals. Pacific Science, 48, 313–324. Entsch, B., Boto, K.G., Sim, R.G. and Wellington, J.T. (1983) Phosphorus and nitrogen in coral reef sediments. Limnology and Oceanography, 28, 465–476. Erez, J. (1990) On the importance of food sources in coral-reef ecosystems, in Ecosystems of the World: Coral Reefs (ed. Z. Dubinsky), Elsevier, Amsterdam, pp. 411–418. Fabricius, K.E., Yahel, G. and Genin, A. (1998) In-situ depletion of phytoplankton by an axoothanthallae soft coral. Limnology and Oceanography, 43, 354–356. Falter, J.L. (1998) Time-variant diagenesis within the redox transition zone of a hydraulically driven sediment. Masters Thesis, University of Hawaii, 101 pp. Falter, J.L. (2002) Mass transfer limits to nutrient uptake by shallow coral reef communities. PhD Dissertation, University of Hawaii, 126 pp. Falter, J.L. and Sansone, F.J. (2000) Hydraulic control of pore water geochemistry within the oxicsuboxic zone of a permeable sediment. Limnology and Oceanography, 45, 550–557. Falter, J.L. and Sansone, F.J. (2000) Shallow pore water sampling in reef sediments. Coral Reefs, 19, 93–97.

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Falter, J.L., Atkinson, M.J. and Langdon, C. (2001) Production–respiration relationships at different time-scales within the Biosphere 2 coral reef biome. Limnology and Oceanography, 46, 1653–1660. Ferrier-Pages, C., Gattuso, J.P., Cauwet, G., Jaubert, J. and Allemand, D. (1998) Release of dissolved organic carbon and nitrogen by the zooxanthellate coral Galaxea fascicularis. Marine Ecology Progress Series, 172, 265–274. Furnas, M., Mitchell, A.W. and Skuza, M. (1995) Nitrogen and phosphorus budgets for the central Great Barrier Reef shelf, Great Barrier Reef Park Authority, 194 pp. Gattuso, J.P., Allemand, D. and Frankignoulle, M. (1999) Photosynthesis and calcification at cellular, organismal, and community levels in coral reefs: a review of interactions and control by carbonate chemistry. American Zoologist, 39, 160–183. Gattuso, J.P., Frankignoulle, M. and Wollast, R. (1998) Carbon and carbonate metabolism in coastal aquatic ecosystems. Annual Review of Ecology and Systematics, 29, 405–434. Gattuso, J.P., Pinchon, M., Delasalle, B. and Frankignoulle, M. (1993) Community metabolism and air–sea CO2 fluxes in a coral reef ecosystem (Moorea, French Polynesia). Marine Ecology Progress Series, 96, 259–267. Gattuso, J.P., Pinchon, M., Delesalle, B., Canon, C. and Frankignoulle, M. (1996) Carbon fluxes in coral reefs. I. Lagrangian measurement of community metabolism and resulting air–sea CO2 disequilibrium. Marine Ecology Progress Series, 145, 109–121. Glynn, P.W. (1973) Ecology of a Caribbean coral reef. The Porites reef-flat biotope: part II. Plankton community with evidence for depletion. Marine Biology, 22, 1–21. Gordon, D.C.J. (1971) Organic carbon budget of Fanning Island Lagoon. Pacific Science, 25, 222–227. Grigg, R.W. (1998) Holocene coral reef accretion in Hawaii: a function of wave exposure and sea level history. Coral Reefs, 17, 263–272. Grigg, R.W., Grossman, E.E., Gittings, S.R., Lott, D. and McDonough, J. (2002) Drowned reefs and antecedent karst topography, Au’au channel, SE Hawaiian Islands. Coral Reefs, 21, 73–82. Haberstroh, P.R. (1994) Wave-forced pore water mixing and nutrient flux in a coral reef framework. PhD Dissertation, University of Hawaii, 249 pp. Haberstroh, P.R. and Sansone, F.J. (1999) Reef framework diagenesis across wave-flushed oxicsuboxic-anoxic transition zones. Coral Reefs, 18, 229–240. Hatcher, B.G. (1997) Organic production and decomposition, in Life and death of coral reefs (ed. C. Birkeland), Chapman & Hall, New York, pp. 140–174. Hearn, C.J., Atkinson, M.J. and Falter, J.L. (2001) A physical derivation of nutrient-uptake rates in coral reefs: effects of roughness and waves. Coral Reefs, 20, 347–356. Hochberg, E.J., Atkinson, M.J. and Andrefouet, S. (2002) Spectral reflectance of coral reef bottomtypes worldwide and implications for coral reef remote sensing. Remote Sensing of Environment (in press). Hoegh-Guldberg, O. and Williamson, J. (1999) Availability of two forms of dissolved nitrogen to the coral Pocillapora damicornis and its symbiotic zooxanthellae. Coral Reefs, 133, 561–570. Huettel, M. and Rusch, A. (2000) Transport and degradation of phytoplankton in permeable sediment. Limnology and Oceanography, 45, 534–549. Huettel, M., Ziebis, W. and Forster, S. (1996) Flow-induced uptake of particulate matter in permeable sediments. Limnology and Oceanography, 41, 309–322. Johannes, R.E., Alberts, J., D’Elia, C., Kinzie, R.A., Pomeroy, L.R., Sottile, W., Wiebe, W., Marsh, J.A.J., Helfrich, P., Maragos, J., Meyer, J., Smith, S.V., Crabtree, D., Roth, A., McCloskey, L.R., Betzer, S., Marshall, N., Pilson, M.E.Q., Telek, G., Clutter, R.I., DuPaul, W.D., Webb, K.L. and Wells, J.M.J. (1972) The metabolism of some coral reef communities: a team study of nutrient and energy fluxes. Bioscience, 22, 541–543. Johannes, R.E. and Gerber, R. (1974) Import and export of net plankton by an Eniwetok coral reef community. Second International Coral Reef Symposium, Brisbane, Australia, 1, 981–988. Johannes, R.E., Wiebe, W.J. and Crossland, C.J. (1983) Three patterns of nutrient flux in a coral reef community. Marine Ecology Progress Series, 12, 131–136.

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Jompa, J. and McCook, L.J. (2002) The effects of nutrients and herbivory on competition between a hard coral (Porites cylindrica) and a brown alga (Lobophora variegata). Limnology and Oceanography, 47, 527–534. Kinsey, D.W. (1985) Metabolism, calcification, and carbon production: I. Systems level studies. Fifth International Coral Reef Congress, Tahiti, 4, 505–526. Kinsey, D.W. and Davies, P.J. (1979) Effects of elevated nitrogen and phosphorus on coral reef growth. Limnology and Oceanography, 24, 935–940. Kinzie III, R.A., Mackenzie, F.T., Smith, S.V. and Stimson, J. (2001) CISNet: linkages between a tropical watershed and reef ecosystems. Project 98-NCERQA, NOAA, 23 pp. Kleypas, J.A., Buddemeier, R.W., Archer, D., Gattuso, J.P., Langdon, C. and Opdyke, B. (1999) Geochemical consequences of increased atmospheric carbon dioxide on coral reefs. Science, 284, 118–120. Koop, K., Booth, D., Broadbents, A., Brodie, J., Bucher, D., Capone, D.G., Coll, J., Dennison, W.C., Erdmann, M., Harrison, P., Hoegh-Guldberg, O., Hutchings, P., Jones, G.B., Larkum, A.W.D., ONeil, J., Steven, A.D.L., Tentor, E., Ward, S., Williamson, J. and Yellowlees, D. (2001) ENCORE: The effect of nutrient enrichment on coral reefs. Synthesis of results and conclusions. Marine Pollution Bulletin, 42, 91–120. Kraines, S., Suzuki, Y., Omori, T., Shitashima, K., Kanahara, S. and Komiyama, H. (1997) Carbonate dynamics of the coral reef system at Bora Bay, Miyako Island. Marine Ecology Progress Series, 156, 1–16. Kraines, S., Suzuki, Y., Yamada, K. and Komiyama, H. (1996) Separating biological and physical changes in dissolved oxygen concentration in a coral reef. Limnology and Oceanography, 41, 1790–1799. Langdon, C. (2001) Review of experimental evidence for effects of carbon dioxide on calcification of reef builders. Abstract. Ninth International Coral Reef Symposium, Bali. Langdon, C., Broecker, W., Hammond, D., Glenn, E., Fitzsimmons, K., Nelson, S.G., Peng, T.H., Hajdas, I. and Bonani, G. (2002) Effect of elevated CO2 on the community metabolism of an experimental coral reef. Global Biogeochemical Cycles, 17(1), 11-1–11-14. Langdon, C., Takahashi, T., Sweeney, C., Chipman, D., Goddard, J., Marubini, F., Aceves, H., Barnett, H. and Atkinson, M.J. (2000) Effect of calcium carbonate saturation state on calcification rate of an experimental coral reef. Global Biogeochemical Cycles, 14, 639–654. Lapointe, B.E. (1997) Nutrient thresholds for bottom-up control of macroalgal blooms on coral reefs in Jamaica and southeast Florida. Limnology and Oceanography, 42, 1119–1131. Larned, S.T. and Atkinson, M.J. (1997) Effects of water velocity on NH4 and PO4 uptake and nutrient-limited growth in the macroalgae Dictyosphaeria cavernosa. Marine Ecology Progress Series, 157, 295–302. Lough, J.M. and Barnes, D.J. (2000) Environmental controls on growth of the massive coral Porites. Journal of Experimental Marine Biology and Ecology, 245, 225–243. Marubini, F. and Atkinson, M.J. (1999) Effects of lowered pH and elevated nitrate on coral calcification. Marine Ecology Progress Series, 188, 117–121. McCook, L.J. (1999) Macroalgae, nutrients, and phase shifts on coral reefs: scientific issues and management consequences for the Great Barrier Reef. Coral Reefs, 18, 357–367. McCook, L.J., Jompa, J. and Diaz-Pulido, G. (2001) Competition between corals and algae on coral reefs: a review of evidence and mechanisms. Coral Reefs, 19, 400–417. Munk, W.H. and Sargent, M.C. (1954) Adjustment of Bikini Atoll to ocean waves. Prof. Paper 260-C, U.S. Geological Survey, pp. 275–280. Naim, O. (1993) Seasonal responses of a fringing reef community to eutrophication (Reunion Island, Western Indian Ocean). Marine Ecology Progress Series, 99, 137–151. Odum, H.T. and Odum, E.P. (1955) Trophic structure and productivity of a windward coral reef community on Eniwetok Atoll. Ecological Monographs, 25, 1415–1444. Pilson, M.E.Q. and Betzer, S.B. (1973) Phosphorus flux across a coral reef. Ecology, 54, 1459–1466. Pomeroy, L.R. and Kuenzler, E.J. (1969) Phosphorus turnover by coral reef animals. Symposium on Radioecology, AEC TID 4500, 3, 474–482.

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Ribes, M., Coma, R. and Gili, J.M. (1998) Heterotrophic feeding by gorgonian corals with symbiotic zooxanthella. Limnology and Oceanography, 43, 1170–1179. Ribes, M., Coma, R., Atkinson, M.J. and Kinzie III, R.A. (2003) Particles removal by coral reef communities: a major source of nitrogen. Marine Ecology Progress Series (accepted). Roberts, H.H., Lugo, A., Carter, B. and Simms, M. (1988) Across-reef flux and shallow subsurface hydrology in modern coral reefs. Sixth International Coral Reef Symposium, Australia, 2, 485–490. Sansone, F.J. (1985) Methane in the reef flat pore waters of Davies Reef, Great Barrier Reef (Australia). Fifth International Coral Reef Congress, Tahiti, 3, 415–419. Sansone, F.J., Tribble, G.W., Andrews, C.A. and Chanton, J.P. (1990) Anaerobic diagenesis within Recent, Pleistocene, and Eocene marine carbonate frameworks. Sedimentology, 37, 997–1009. Sargent, M.A. and Austin, T.S. (1949) Organic productivity of an atoll. Transactions of the American Geophysical Union, 30, 245–249. Schlichter, D. and Liebezeit, G. (1991) The natural release of amino acids from the symbiotic coral Heteroxenia fuscescens (Ehrb.) as a function of photosynthesis. Journal of Experimental Marine Biology and Ecology, 150, 83–90. Sebens, K.P., Grace, S.P., Helmuth, B., Maney Jr., E.J. and Miles, J.S. (1997) Water flow and prey capture by three scleractinian corals, Madracis mirabilis, Montastrea cavernosa, and Porites porites in a field enclosure. Marine Biology, 131, 347–360. Seitzenger, S.P. and D’Elia, C.F. (1985) Preliminary studies of denitrification on a coral reef. Symposium for Undersea Resarch, NOAA’s Undersea Research Program, 3, 199–208. Simkiss, K. (1964) The inhibitory effects of some metabolites on the precipitation of calcium carbonate from artificial and natural sea water. Journal du Conseil International pour l’Exploration de la Mer, 29, 6–18. Skyring, G.W. and Chambers, L.A. (1976) Biological sulfate reduction in carbonate sediments of a coral reef. Australian Journal of Marine and Freshwater Research, 27, 595–602. Smith, S.V. (1984) Phosphorus versus nitrogen limitation in the marine environment. Limnology and Oceanography, 29, 1149–1160. Smith, S.V. and Buddemeier, R.W. (1992) Global change in coral reef ecosystems. Annual Review of Ecology and Systematics, 23, 89–118. Smith, S.V. and Marsh, J.A. (1973) Organic carbon production on the windward reef flat of Eniwetok Atoll. Limnology and Oceanography, 18, 953–961. Sorokin, Y.I. (1991) Biomass, metabolic rates and feeding of some common reef zoantharians and octocorals. Australian Journal of Marine and Freshwater Research, 42, 729–741. Steven, A.D.L. and Atkinson, M.J. (2003) Nutrient uptake by coral-reef microatolls. Coral Reefs (in press). Stoddart, D.R. (1969) Ecology and morphology of recent coral reefs. Biological Review, 44, 433–498. Suzumura, M., Miyajima, T., Hata, H., Umezuwa, Y., Kayanne, H. and Koike, I. (2002) Cycling of phosphorus maintains the production of microphytobenthic communities in carbonate sediments of a coral reef. Limnology and Oceanography, 47, 771–781. Szmant, A. and Forrester, A. (1996) Water column and sediment nitrogen and phosphorus distribution patterns in the Florida Keys, USA. Coral Reefs, 15, 21–41. Tarrant, A. (2002) Estrogen action in scleractinian corals: sources, metabolism and physiological effects. PhD Dissertation, University of Hawaii, 186 pp. Thomas, F.I.M. and Atkinson, M.J. (1997) Ammonium uptake by coral reefs: Effects of water velocity and surface roughness on mass transfer. Limnology and Oceanography, 42, 81–88. Tribble, G.W., Sansone, F.J. and Smith, S.V. (1990) Stoichiometric modeling of carbon diagenesis within a coral reef framework. Geochimica et Cosmochimica Acta, 54, 2439–2449. Tribble, G.W., Sansone, F.J., Buddemeier, R.W. and Li, Y.-H. (1992) Hydraulic exchange between a coral reef and surface sea water. Geological Society of America Bulletin, 104, 1280–1291.

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Ward, S. and Harrison, P. (2000) Changes in gametogenesis and fecundity of acroporid corals that were exposed to elevated nitrogen and phosphorus during the ENCORE experiment. Journal of Experimental Marine Biology and Ecology, 246, 179–221. Webb, K.L., Paul, W.D.D., Wiebe, W., Sottile, W. and Johannes, R.E. (1975) Enewetak (Eniwetok) Atoll: aspects of the nitrogen cycle on a coral reef. Limnology and Oceanography, 20, 198–209. Wiebe, W.J., Johannes, R.E. and Webb, K.L. (1975) Nitrogen fixation in a coral reef community. Science, 188, 257–259. Wilkinson, C.R., Williams, D., Sammarco, P.W., Hogg, R.W. and Trott, L.A. (1984) Rates of nitrogen fixation on coral reefs across the continental shelf of the central Great Barrier Reef. Marine Biology, 80, 255–262. Williams, S.L. and Carpenter, R.C. (1998) Effects of unidirectional and oscillatory water flow on nitrogen fixation (acetylene reduction) in coral reef algal turfs, Kaneohe Bay, Hawaii. Journal of Experimental Marine Biology and Ecology, 226, 293–316. Yahel, G., Post, A.F., Fabricius, K., Marie, D., Vaulot, D. and Genin, A. (1998) Phytoplankton distribution and grazing near coral reefs. Limnology and Oceanography, 43, 551–563. Yahel, G., Sharp, J.H., Marie, D., Hase, C. and Genin, A. (2003) In-situ feeding and element removal in the symbiotic-bearing sponge Theonella swinhoei: bulk DOC is the major source for carbon. Limnology and Oceanography (in press).

3

Fjords J.M. Skei, B. McKee and B. Sundby

3.1

Introduction

A fjord is a type of estuary found in parts of the world that were previously glaciated. Fjords interface and buffer the zone between land and sea, and – by volume, if not by area – constitute an important part of the coastal zone. The total volume of water contained in the fjords of the world is similar to the total volume in freshwater lakes (about 1.4 × 1014 m3, Syvitski et al., 1987). Fjords are the deepest of all estuaries, and fjords of more than 1000 m depth are found in Chile and Norway. Some fjords have sills at the mouth, others do not. Where tidal and wind mixing is weak, the rate of deep water exchange in silled fjords tends to be slow, and permanent or temporary stagnant conditions may develop in the deep water (Skei, 1988a). The variety of environmental conditions found in fjords is both a challenge and an opportunity to geochemists (Fig. 3.1). For example, the interface between oxic and anoxic conditions may be located within the sediment column, as is the case of typical shelf sediments, within the water column itself, or even within the euphotic zone (Skei, 1982). Fjords are therefore well suited to study redox processes associated with sharp redox gradients (Skei, 1983). The highly variable input of fresh water and sediment affects circulation and sedimentation patterns both within and between fjords. High sedimentation rates where the sediments enter a fjord, create deposits of poorly consolidated sediments inclined to slumping under their own weight, or when disturbed physically. Slumping events redistribute these sediments to deeper areas of the fjord (Fig. 3.1). Fjord sediments, therefore, often contain layers of different composition and different properties that have accumulated at different rates. Seasonal variability in sediment and of organic matter input, as well as the occasional sediment redistribution by slumping, play an important role in biogeochemical processes. The spatially and temporally variable sedimentation patterns that are typical in a fjord can easily confound the interpretation of historical records of both particle bound and post-depositionally mobile sediment components. Hence, a good understanding of both sedimentation history and sediment diagenesis is essential to understand the sedimentary record. Some fjords are located in arctic and sub-arctic regions where the influence of man is minimal; others are located near population and industrial centres and can be heavily contaminated (Skei et al., 2000). This spectrum of conditions

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Light Sediment Fresh C water

Primary production Sed plume Stratification

Ice Tidal energy Sill

Mixed layer

Resuspension

Slumping

Basin O2

Episodic flushing

Sediment Nutrients

Upwelling C

Oxic

Anoxic

Benthos

Redox

Burial

Fig. 3.1 Interrelationship of physical, geological, biological and chemical processes in fjords (after Syvitski et al., 1987).

offers the opportunity to examine the functioning of fjord systems both in the presence and absence of human influence. Fjord research has a long history, but most of the focus has been on physical oceanography, sedimentology and biology, and comparatively less effort has been made to study chemical and geochemical aspects of fjords (Holtedahl, 1967; Freeland et al., 1979; Syvitski & Skei, 1983; Syvitski, 1984; Syvitski et al., 1987). The objective of this chapter is to review the present knowledge of the biogeochemistry of fjords, emphasizing sediment diagenesis and element cycling in oxic and anoxic fjords. 3.1.1

Definition and origin of fjords

There are several definitions of fjords. The most common definition is ‘a fjord is a deep, high-latitude estuary which has been (or is presently being) excavated or modified by land-based ice’ (Syvitski et al., 1987). This definition implies that fjords are only found in the northernmost and southernmost regions of the earth that once were ice covered. As a geological feature, fjords are the youngest of all estuaries since they were formed during the retreat of the glaciers and

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the fluctuations of the sea level that have taken place since the last glacial maximum some 17 000 years BP. Fjords can be classified as immature, nonsteady state systems. Some of the earliest studies of fjords focused on their origin (Gregory, 1913). The characteristic geomorphology of fjords generated a great deal of speculation, ranging from a purely tectonic origin to fluvially or glacially modified mountain valleys. The accepted view today is one of accommodating compromise, i.e. fluvial action along fault lines, with subsequent and dominant excavation by glaciers following the path of least resistance. 3.1.2

The public and scientific interest in fjords

Fjords are diverse with respect to topography, hydrography, current regimes, tidal influence, chemistry, sediment input and transport, oxygen conditions, trophic status, productivity, human impact, commercial utilization, etc. This implies that we need to understand how fjords work as systems taking into account their natural differences as well as anthropogenic impacts. One of the scientific interests of fjords is related to the relative simplicity of the fjord system compared to other types of estuaries, which make them more amenable to modelling. Fjords are often long and narrow with a fresh water input at the head and a sill at the mouth. This creates a predicable water circulation pattern whereby the deep water is only exchanged periodically. The periodicity depends on the time required for the rate of vertical mixing to decrease the density of the bottom water to values below the density of the ocean water residing outside the mouth of the fjord. Where the sill is shallow and the mixing rate weak, the water residing below the sill depth will not be exchanged (flushed) very frequently, and between the flushing events, the deep water may turn stagnant and even anoxic. These are ideal conditions for studying how the chemistry of estuarine water evolves over time. Such fjords have been used for kinetic studies to determine the rate constants of chemical reactions (Shapiro et al., 1997). A deep water exchange event, which replaces the stagnant water with oxygenated water, signals the beginning of a new period of stagnation and oxygen depletion. This property makes some fjord basins attractive, as a means to understand such important phenomena as oxygen depletion, denitrification, sulfate reduction (Cutter & Kluckhohn, 1999) and methanogenesis (Lidstrom, 1983). The long residence time of the bottom water also makes fjords suitable for examining the importance of sediment–water interactions (Sternbeck et al., 1999), including the release of re-mineralized nutrients to the bottom water, the release of contaminants from sediments, and the exchange of gases. Silled fjords are not necessarily stagnant. Where tidal energy and winds are sufficient to provide efficient and frequent vertical mixing, the bottom water is exchanged frequently and its residence time is too short for the oxygen content

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to be completely depleted. In fact, the bottom water in most of the fjords of the world is oxic, and only rarely is the basin water anoxic over long periods. Productivity in fjords can be highly variable, depending on the availability of nutrients and the residence time of the surface water. Sometimes algal blooms occurring in fjords can be so intense that the water is discoloured. Blooms of certain species of algae create economic problems for the aquaculture industry and sometimes health problems for the public when the algae are toxic. Thus fjords represent both eutrophic and oligotrophic systems. The decay of the remains of algal blooms increases the oxygen consumption in the underlying water. Depending on the water depth, the flux of organic matter to the sediment may be high, creating anoxic conditions near the sediment–water interface. This has important consequences for the fate of contaminants in sediments (Skei et al., 2000), as sulfides in the sediments may bind and immobilize certain toxic trace metals. Fjords offer a great variety and spatial gradients of a number of interesting features, including water depth, sill depth, redox conditions, salinity, climate zones, level of anthropogenic influence, freshwater and sediment input, stratification, sediment types, sediment ecology, sediment biogeochemistry, and biological productivity. All this may be studied within relatively small geographical distances using small ships, which is why fjord research can yield high scientific returns with modest investments in cost and effort (Skei, 1983). One has to understand how fjords work as natural systems in order to utilize this valuable resource in a sustainable manner. The physical transport of water and sediments determines the recipient capacity with respect to human influence. In the past, the pressure from hydroelectric power development, which influences the fresh water input, the circulation and exchange regime, and the biology of fjords, was intense. Today, there are new threats from the aquaculture industry, which is rapidly expanding in fjords. Aquaculture activities release increasing quantities of organic matter and nutrients, which may promote eutrophication and ultimately lead to oxygen depletion. This situation is a classical example of the inevitable conflicts between concerns for environmental quality and the need for economic growth in regions where alternative commercial activities are scarce. Balancing these apparently conflicting demands will require the knowledge and understanding of the fjord environment that scientists can provide. In addition to fish farming, which is a relatively recent phenomenon, hydroelectric power dependent chemical industry has been using fjords as recipient for waste for more than 50 years, particularly in Scandinavia and Canada. Smelters located at the head of fjords have contaminated water, sediments, and organisms with polycyclic aromatic hydrocarbons (PAHs) and heavy metals (Naes et al., 1995). The level of contamination in fish and shellfish exceeds safety levels established by the World Health Organization in 15 fjord locations in Norway, and their use for human consumption has been restricted. These

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restrictions are partly due to elevated levels of PCB in fish liver. The scientific challenge is to discover if the contamination is due to primary discharges of PCB from land or contaminated fjord sediments. 3.2

Sediment diagenesis in oxic fjords

Diagenesis is defined as the sum of all the changes that take place in sediments accumulating under water from the moment the sediment is laid down until it is finally converted to sedimentary rock. The earliest stages of diagenesis are characterized by strong interactions between the solid and fluid components of the sediment, and between the sediment and the overlying water column. The spatially and temporally variable sedimentation patterns in fjords naturally influence diagenesis, and a good knowledge of the sedimentation history is essential to understand how diagenesis affects fjord sediments. The following discussion will focus on the Saguenay Fjord, where such information is available. 3.2.1

The Saguenay Fjord

The Saguenay Fjord is a 93 km long, 1–6 km wide submarine valley in eastern Canada. It is composed of several basins, the largest of which is 275 m deep. The basins are separated by shallow sills. The shallowest of these (20 m) separates the fjord from the St. Lawrence Estuary. The fjord connects the lower St. Lawrence Estuary to a 78 000 km2 basin draining the Precambrian rocks of the Canadian Shield. However, uplifted post-glacial marine clays (Laflamme Sea clays) are the principal source of sediments to the fjord (Schafer et al., 1983, and references therein). The region near the head of the fjord is heavily industrialized and, in the early 1970s, it was discovered that the fjord sediments contained anomalously high levels of mercury (Loring, 1975). This discovery gave impetus to a series of studies that have provided detailed data on the sedimentation regime. Although originally intended to help elucidating the history of mercury contamination (Smith & Loring, 1981), these studies also served to document the recent climatology of the region (Smith & Schafer, 1987; Smith, 1988). 3.2.2

Sedimentation

Sediment accumulation rates in the Saguenay Fjord decrease with increasing water depth and with distance from the mouth of the Saguenay River. The rates range from about 7 cm y−1 (4 g cm−2 y−1) near the head of the fjord to 0.1 cm y−1 (0.07 g cm−2 y−1) in the centre of the 275 m deep main basin (Smith & Walton, 1980). This study, which was based on the Pb-210 dating method, also revealed a slumping event in one of the cores from the head of the fjord. The

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timing of the event was determined as 1972 ± 1 year, which coincides with a catastrophic landslide that occurred in May 1971 at Saint Jean Vianney on the banks of a tributary to the Saguenay River. This landslide released 25 × 106 tons of postglacial marine sediment into the tributary. The Saint Jean Vianney landslide is not unique in the history of the Saguenay Fjord. In 1663, perhaps the largest earthquake ever witnessed in eastern North America struck the Saguenay Basin, triggering submarine slides over a wide region. Over 3 km3 of Holocene sediments collapsed from the margins of the fjord and were redeposited on the fjord bottom. The new deposits reached thicknesses of 100 m and extended over an area of 100 km2 (Syvitski & Schafer, 1996). In July of 1996, a catastrophic flood in the Saguenay headlands eroded 9–15 × 106 tons of post-glacial sediments and carried them into the fjord. At the head of the fjord, several tens of centimeters of sandy clays were rapidly deposited over the indigenous sediment (Pelletier et al., 1999). 3.2.3

Composition of the rapidly deposited layers

The rapidly deposited layers in Saguenay Fjord sediments can be distinguished from the sediments that accumulate under normal circumstances (indigenous sediments) by their lower content of organic carbon (<1%) and higher content of carbonate (>2% as CaCO3) (Schafer et al., 1983; Mucci & Edenborn, 1992). The isotopic signature of the organic carbon in these layers (δ13C = −29‰) corresponds to that of Laflamme Sea sediments (St-Onge & Hillaire-Marcel, 2001). Compared to the rapidly deposited layers, the indigenous sediments in the fjord contain up to 2.5% organic carbon and less than 0.5% CaCO3, and the isotopic signature of the organic carbon tends to be heavier (δ13C = −26‰) (St-Onge & Hillaire-Marcel, 2001). The higher organic carbon content reflects the addition of organic matter from a variety of sources, including terrestrial plant material and aquatic organic matter produced by freshwater and marine organisms. At one time, the release of organic matter from pulp and paper industries was so high that it influenced both the concentration and the composition of organic matter in sediments (Schafer et al., 1983; Louchouarn et al., 1997; Louchouarn & Lucotte, 1998). The lower CaCO3 content and the absence of calcareous foraminifers in the indigenous sediments have been ascribed to dissolution of fossil carbonate during transport (Schafer et al., 1980) and during the initial stages of burial (Mucci et al., 2000b). 3.2.4

Sulfate reduction and sulfide accumulation

Sulfate reduction rates in the deep basin sediments of the Saguenay Fjord are relatively low in comparison to other coastal environments (32 nmol cm−2 d−1 integrated over the top 30 cm), and the pore water is only weakly depleted in sulfate. Sulfate reduction is more intense, however, in the rapidly accumulating

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sediments at the head of the fjord (110 nmol cm−2 d−1), and the pore water is strongly depleted in sulfate (Edenborn et al., 1987). The sulfate reduction rate is highly correlated with the organic carbon content and is negligible within the rapidly deposited organic-carbon poor landslide layer (Fig. 3.2). Sulfate reduction in organic rich coastal marine sediments usually leads to the precipitation of pyrite (FeS2). Among several possible reactions that may lead to pyrite, the conversion of iron monosulfides (FeSx where 0.9< × <1.5) to pyrite in the presence of zero-valent sulfur species has been well documented for estuarine and marine sediments. Because they are converted to pyrite, iron monosulfides are normally found as a minor constituent of the pool of reduced sulfur species, and in most environments studied, the FeSx–sulfur:FeS2–sulfur ratio is much less than one. The Saguenay Fjord sediments are anomalous in this regard. The degree of pyritization, a measure of the conversion of reactive iron to pyrite, is low throughout the fjord, and the FeSx–S:FeS2–S ratio (where monosulfide–sulfur is represented by Acid Volatile Sulfide, or AVS) exceeds 1 throughout most of the indigenous anoxic sediments. In the sediments at the head of the fjord, where the rate of sulfate reduction is high, the FeSx–S:FeS2–S ratio reaches values as high as 7 (Gagnon et al., 1995). The conditions required for pyrite formation (reduction of sulfate and availability of reactive iron phases) exist in the Saguenay Fjord. Indeed, the high

0

Hg (ppm)

0

C (%)

400 0

12

100

% of landslide material

0

AVS and pyrite sulfur (μmol g–1)

80

0 A

Depth (cm)

Hg

SRR

B

C

C

10

20 0 Pyrite-S

AVS

30 0

2

4

Sulfate reduction rate (SRR) (pmol cm–3 h–1)

Fig. 3.2 (a) Depth distribution of total Hg, %C and bacterial sulfate reduction rates (SRR). (b) Sediment mixing curves determined across the landslide layer based on Hg and %C data presented in (a). (c) AVS and pyrite sulfur concentrations at station S-1, October 1984. Dashed lines indicate approximate location of lower and upper boundaries of the landslide.

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degree of sulfidization (a measure of the conversion of reactive iron to iron monosulfides) in the sediments at the head of the fjord shows that pyrite formation is not limited by the availability of reactive iron. Gagnon et al. (1995) concluded that the conversion of AVS to pyrite in these sediments is inhibited by the efficient precipitation of sulfide by abundant reactive iron. This prevents the buildup of reduced sulfur in the sediment pore water and limits the flux of H2S to the oxidizing zones of the sediment. Therefore, the formation of sulfur species of intermediate oxidation states such as elemental sulfur and polysulfides, required for the conversion of monosulfides to pyrite, is severely restricted. This study was followed up with an examination of the concentration and distribution of reduced sulfur species in the pore water by Gagnon et al. (1996a). They found that the concentration of dissolved sulfides and polysulfides in the top 40 cm of the sediment was low and usually below 1 μM. They also documented the distributions of dissolved organic sulfides and sulfur species resulting from the oxidation of sulfide, including thiosulfate, sulfite, and elemental sulfur. With the exception of thiosulfate, the concentration of each of these species was generally higher than the concentration of dissolved sulfides. Thus, the concentration of elemental sulfur ranged from 5 to 20 μM; the highest values occurred in the sediment at the head of the fjord, near the upper boundary of the sediment layer deposited during the 1971 landslide event. Coinciding with these high values of elemental sulfur were high concentrations (10–35 μM) of dissolved organic sulfur. Thus, although reactive iron immobilizes most of the sulfide produced during sulfate reduction, a small fraction is nevertheless oxidized to intermediate sulfur species. The pathways of the many reactions involving sulfur and iron species, adapted to the situation in the Saguenay Fjord, are illustrated in Fig. 3.3. SO

4 2–

M re icro du b ct ial io n

n atio oxid

H2S

chemical oxidation

SO32– FeS S2O32–

Fe2+ RSR

oxidation

FeS

S0

HS–

Sn2–

FeS

FeS FeS2

Fig. 3.3 Proposed reaction scheme of sulfur in the Saguenay Fjord sediments. The width of the arrows provides a qualitative indication of the relative importance of each step.

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3.2.5

73

Mercury diagenesis

The early work on mercury in the Saguenay Fjord, which documented the abundance of mercury in sediments and the history of mercury pollution in the fjord (Loring, 1975; Smith & Loring, 1981), ultimately led to the closure of the plant that was responsible for most of the Hg-discharge and the introduction of new environmental regulations. During the time the plant was in operation, it consumed more than 300 tons of Hg of which an estimated 60 tons is now buried in the fjord sediments (Loring & Bewers, 1978). Considering the potential for post-depositional remobilization, recent studies have focused on the diagenesis of mercury. Despite the sharp reduction in mercury discharge that followed the plant closure in 1976, the mercury content of surface sediments has remained high (500 ng g−1) when compared to pre-industrial Hg levels (<100 ng g−1). This raised the possibility that the surface enrichment might be due to upward migration of Hg from deeper, contaminated sediment layers. Gagnon et al. (1997) found variable concentrations of dissolved Hg in pore water, ranging from 17 to 500 ng g−1. No obvious relationship could be found between dissolved and solid-phase Hg concentrations. Methyl mercury was present in the pore water (up to 10 ng Hg l−1), which accounts for it being less than 1% of the total CH3Hg(II) in the sediment. The flux of dissolved Hg from the sediment to the overlying water, estimated from the concentration gradient across the sediment– water interface, ranged from 9 to 36 ng cm−2 y−1, more than one order of magnitude lower, than fluxes from contaminated sediments underlying anoxic or almost anoxic bottom water (Bothner et al., 1980). It was suggested that the presence of oxidizing surface sediments inhibits the transfer of Hg to the water column, possibly because of the scavenging ability of iron and manganese oxides. As a proportion of the total flux of dissolved Hg, the flux of methyl mercury is insignificant (0.2%) (Gagnon et al., 1997). The diffusive flux of Hg from the deeper, anoxic layers is insufficient to explain the enrichment of the sediment surface layer. Gobeil and Cossa (1993) estimated that this flux amounts to no more than 2.5% of the mercury transported to the sediment on settling particles. An analysis of suspended particulate matter in the fjord confirmed this, showing Hg concentrations in suspended particulate matter (SPM) similar to those in the surface sediment (Gagnon et al., 1996b). These studies effectively eliminate diagenesis as the reason for Hg enrichment in the surface sediments. 3.2.6

Phosphorus and arsenic geochemistry

The similarities and the differences in geochemical properties of phosphorus and arsenic are reflected in the way these elements respond to diagenesis in the Saguenay Fjord (Mucci et al., 2000a). Both the elements are supplied to the sediment on settling particles, which appear to scavenge both P and As from

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the water column as they settle to the sea floor. Evidence for scavenging is strong: the P and As content of particulate matter increases with depth, and the distribution of phosphate in the fjord water is non-conservative. The concentration of either element in Saguenay River water is lower than in seawater, which may be their main source. Industrial sources of arsenic exist, but their quantitative significance is not known. Compared to the anoxic sediment, the oxic surface sediment is enriched in As (up to 345 nmol g−1). The enrichment is weak where the sedimentation rate is high, but enrichment becomes stronger with decreasing sedimentation rate. Phosphorus enrichment in the surface layer is evident only where the sedimentation rate is low. However, upon burial, both P and As are released from the solid phase to the pore water, irrespective of the sedimentation rate, and the distributions of both elements show subsurface maxima that correspond to the depth where the dissolution rates of solid phases of P and As are maximum. From the relative location of these maxima, it can be concluded that As is released to the pore water at shallower depths than at which phosphate is released. Mucci et al. (2000a) proposed that As is released upon the reduction of As(V) to soluble As(III), and that this takes place earlier than the reduction of Fe(III) to Fe(II) in the diagenetic sequence. In contrast, the

Water column Sequestered phosphorus

Adsorbed phosphorus

Dissolved phosphorus

Organic phosphorus

Sequestered phosphorus

Adsorbed phosphorus

Dissolved phosphate

Organic phosphorus

Adsorbed phosphorus

Dissolved phosphate

Organic phosphorus

Sediment

Fe(II) oxidation Fe(III) reduction

Sequestered phosphorus

Fig. 3.4 The pathways of reactions involving phosphorus in iron-rich sediments.

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release of P takes place upon the reductive dissolution of the iron hydroxide carrier phase. The release of P and As to the pore water creates upward directed concentration gradients and drives a flux of both the elements towards the sediment surface. Phosphate will be readsorbed by authigenic iron oxides and will continue to cycle across the Fe(III)/Fe(II) redox boundary (Sundby et al., 1992). Figure 3.4 illustrates the pathways of reactions involving phosphorus, in iron rich sediments. Arsenite, however, appears to migrate well into the oxic surface layer, and it is suggested that it may be oxidized by authigenic manganese oxides before being adsorbed by iron oxyhydroxides present at the same depth (Mucci et al., 2000a). Resuspending anoxic Saguenay Fjord sediment in aerated seawater and monitoring the evolution of dissolved Fe, Mn, and As species, gave results that are consistent with faster oxidation kinetics of As(III) to As(V) in the presence of manganese oxides (Saulnier & Mucci, 2000). The possibility that manganese oxides catalyze arsenite oxidation is intriguing because arsenic retention by sediments would then depend on the availability of manganese oxides and the thickness of the oxic layer, as well as on co-precipitation with iron sulfides at depth (Mucci et al., 2000a). 3.2.7

Non-steady-state diagenesis

Early diagenesis is often described as a steady-state phenomenon whereas, in reality, the sediment composition can vary temporally for a variety of reasons including non-steady sedimentation events such as those occurring in the Saguenay Fjord. The significance of such events for diagenesis of the fjord sediments was first discussed by Mucci and Edenborn (1992) who pointed out that the 1971 landslide layer influenced the migration and deposition patterns of mobilizable elements, acting alternatively as a sink and a source of reduced Fe and Mn, as the redox boundary migrated through it (Fig. 3.2). A catastrophic flash flood in July 1996, when as much as 50 cm of postglacial sediments were deposited in less than two days in the upper reaches of the Saguenay Fjord, provided an opportunity to study non-steady-state diagenesis in more detail (Deflandre et al., 2002). A set of stations were visited immediately after the flood and once every year thereafter to document the disruption of the existing steady state and the evolution of the sediment towards a new steady state. The deposition of the flood material modified the distribution patterns of reactive Mn and Fe via the reduction of Mn and Fe oxides delivered with the flood material as well as the reduction of oxides present at the former sediment–water interface, now buried under the flood layer. Most of the dissolved manganese migrated to the new sediment–water interface where a new Mn-rich layer soon began to form. Unlike manganese, the dissolved Fe(II) did not migrate. It was precipitated as sulfides and trapped at or close to

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the old interface. These results are in agreement with the conceptual model proposed by Mucci and Edenborn (1992). The rapid deposition of the flood material revealed diagenetic reactions that tend to be imperceptible under steady-state conditions. The temporary presence of manganese oxides at the former sediment–water interface, now anoxic, appeared to react with ammonia to produce nitrate anaerobically. This is consistent with recent laboratory and field studies that have related anaerobic nitrate production to manganese oxide reduction (Hulth et al., 1999; Anschutz et al., 2000). Furthermore, tight correlations were observed between dissolved organic carbon and both dissolved Fe(II) and Mn(II) in the sediment deposited during the flood. This may reflect the release of adsorbed organic matter to the pore waters as metal oxides are being reduced. 3.3

Elemental cycling in anoxic waters

Diagenesis and elemental cycling in oxic or suboxic sediments in Saguenay Fjord is to some extent governed by different processes than in fully anoxic systems, where organic matter oxidation and sulfate reduction have been taking place without interruption for a considerable length of time. This has led to a situation where sulfide is present in the water column. 3.3.1

Chemical tracers

Studies of chemical tracers with known input sources have provided important insights into the mixing dynamics and water exchange in highly stratified anoxic fjords. These studies also yield valuable information about the stability and biogeochemistry of the tracers (such as dissolved gases, CFCs, CCl4, tritium and helium, stable isotopes and anthropogenic radioisotopes) under anoxic conditions. Measurements of anthropogenic halocarbons CCl3F (CFC-11) and CCl2F2 (CFC-12) made in Saanich Inlet (Bullister & Bing-Sun, 1995) exhibit a pronounced depletion of dissolved CFC-11 in the subsurface anoxic zones. These zones are strongly reducing and are characterized by the presence of hydrogen sulfide (H2S). Models incorporating the atmospheric input histories of these CFCs and observed distributions were used to estimate residence times for water in these zones and first-order in situ removal rates for CFC-11. Vertical distributions of CFC-12, CFC-11, and CCl4 (carbon tetrachloride) in Saanich Inlet were used to estimate the renewal rates of bottom water, and to study the non-conservative behavior of these compounds (Lee et al., 1999). Dissolved CFC-11 and CCl4 were strongly depleted in the deep anoxic waters, with CFC-11 concentrations <30% of saturation levels, and CCl4 concentrations near blank (<1% of saturation) levels. In contrast, CFC-12 was near saturation throughout the water column. Incubation experiments in Saanich Inlet showed rapid removal of CCl4 in anoxic waters but undetectable removal rates for

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CFC-11 and CFC-12 over a 10 day period. Over periods of months to years, CFC-12 behaved conservatively in these studies while CFC-11 was removed slowly with first-order removal coefficients ranging from ~0.5 y−1 to 1.2 y−1 over a range of temperatures. In oxygenated waters, CCl4 was reactive with a removal rate constant of ~0.69 y−1 while CFC-11 and CFC-12 was unreactive (Lee et al., 1999). In Framvaren Fjord (Norway), turbulent vertical mixing appears to be the primary mechanism in which surface waters are transported deeper into the water column, below the O2/H2S interface. Shapiro et al. (1997) determined the rate of turbulent mixing by fitting a 1D model to the vertical distributions of helium and tritium. Comparisons of modelled and observed CFC profiles were then used to obtain first-order degradation rates for CFC-11 and CFC-12. In the anoxic bottom waters, CFC-11 was degraded at a rate of 6–9 y−1, and CFC-12 was found to have a maximum degradation rate of 0.01–0.03 y−1. Sanchez et al. (1994) used plutonium as a tracer for water exchange between these Framvaren Fjord and the Skagerrak Sea outside. The 238Pu/239,240Pu ratio from fuel reprocessing facilities in UK and France is found only in the surface layer of the fjord (upper 20 m) while global fallout Pu ratios are observed in the deeper, permanently anoxic zone of this fjord. Plutonium oxidation state measurements show that water exchange occurs between Framvaren and the basin outside by a process of interleaving, with subsequent reduction of oxidized Pu species as they reach the anoxic zone. The effect of this interleaving appears to be minimal for the bottom waters of Framvaren (below 90 m) because the 14C age of the total dissolved inorganic carbon (Ct) in the anoxic bottom waters is about 1600 years indicating a bioproduction in the period 8000 years BP to AD 1853 when a channel was opened between the fjord outside (Helvikfjord) and Framvaren (Dyrssen et al., 1996). 3.3.2

Cycling of carbon and nutrients

During the past decade, stable C and N isotopes have been used very effectively in anoxic fjords to improve our understanding of carbon and nitrogen cycling in these environments. Stable isotope analysis (δ15N and δ13C) of dissolved inorganic and particulate organic matter (POM) was used in Framvaren Fjord to evaluate the cycling of carbon and nitrogen within the water column, and deposition to the sediments. The carbon isotopic distribution of dissolved inorganic carbon (DIC) and particulate organic carbon (POC) within the interface suggests that the distinct microbial flora (Chromatium sp. and Chlorobium sp.) fractionate inorganic carbon to different degrees (Velinsky & Fogel, 1999). Nitrogen isotopic mass balance calculations suggest that approximately 75% of the material sinking to bottom sediments was derived from the dense particulate maximum near the O2/H2S interface (Velinsky & Fogel, 1999). The 13C signature in the Framvaren bottom waters (19 parts per thousand) is

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mainly determined by decomposition processes that are acting upon the organic matter produced in the surface waters (Dyrssen, 1999). Isotopic fractionation of dissolved ammonium during consumption by bacteria (using 15N) was found to be greatest near the O2/H2S interface for Saanich Inlet and Framvaren Fjord (Velinsky et al., 1991). Isotope enrichments were between 20 and 30 parts per thousand in Framvaren Fjord. Biosynthetic uptake of NH4+ rather than nitrification was responsible for the fractionation in Framvaren, while the isotopic fractionation in Saanich Inlet appears related to in situ regeneration of NH4+ with little isotopic fractionation between dissolved and particulate nitrogen. The high δ15N-NH4+ at the O2/H2S interface, which is a consequence of microbial uptake, results in extremely depleted δ15N – particulate nitrogen of approximately 1 part per thousand within the particulate maximum (Velinsky & Fogel, 1999). The relation between silica, ammonium and phosphate in Framvaren can be best understood if part of the ammonium is being removed by denitrification (Dyrssen, 1999). The distribution of nitrogen isotopes in Framvaren sediments varied from 2 parts per thousand at the surface to approximately 6 parts per thousand at 30 cm. The nitrogen isotopic variations with depth may be an indicator of the depth or position of the O2/H2S interface in the fjord. Low sediment δ15N indicates that the interface was within the photic zone of the water column, while more enriched values suggested that the interface was lower in the water column, potentially allowing for less fractionation during biological incorporation of dissolved inorganic nitrogen. Results indicate that the dense layers of photo-autotrophic bacteria in the upper water column impart unique carbon and nitrogen isotopic signals that aid in tracking processes within the water column and the deposition to the sediments (Velinsky & Fogel, 1999). 3.3.3

Trace element and radionuclide cycling

The cycling of trace elements and radionuclides in natural environments is closely coupled to the cycling of Fe, Mn and C, because metal oxides and organic carbon are usually dominant carrier phases (Fig. 3.5). This strong coupling has been especially noted by several studies in anoxic fjords. The temporal and spatial gradients in the particulate Mn distribution in Saanich Inlet were used to calculate the in situ rate of Mn precipitation (Grill, 1982). Comparisons of calculated and theoretical rates suggest that two parallel processes control the precipitation of Mn in this environment: bacterial oxidation and an inorganic autocatalytic oxidation reaction. Rare earth element (REE) and manganese distributions in Saanich Inlet reveal that redox cycling of Ce and Mn takes place across an oxic/suboxic boundary approximately 20 m shallower than the O2/H2S interface; the remaining, strictly trivalent REE undergo scavenging above this oxic/suboxic boundary, and regeneration from dissolving host phases below it (German & Elderfield, 1989). These processes

2– 3

0

S

SO42–-reducing bacteria

S-reducing bacteria

Green S bacteria

Purple S bacteria

H2S

O2

C

C

D

B

A

E

Mn(II)

Mn(IV)

Mn(IV)

Mn(II)

by oxidation of Mn(II)

Production of Mn(IV)

Mn CYCLE

Reduction

Mn

Aerobic

E

H, I

F, G

U(VI)

C

C

U(VI)

U(IV)

URANIUM CYCLE

A n o x i c

O x i c

Fig. 3.5 The cycling of sulfur, manganese and uranium at the oxic–anoxic interface.

[A] Particle settling; [B] Reduction of Mn(IV); [C] Diffusion/advection; [D] Oxidation of Mn(II); [E] Authigenic mineral formation; [F] Uranium redox reactions; [G] U:Mn (carrier phase) reactions; [H] U:microbial interactions; [I] U:DOC associations.

S2O32–

SO42– SO

2– 3 S2O32–

SO42– SO

S0

Chemical oxidation

Aerobic S bacteria

S CYCLE

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result in pronounced removal of Ce from solution above the oxic/suboxic boundary due to oxidative precipitation along with preferential uptake of the lighter trivalent REE during scavenging. Swarzenski et al. (1999a) noted that a sharp decrease in pH above the O2/H2S interface in Framvaren Fjord facilitates the aerobic dissolution of MnO2. In contrast, Fe(II) concentrations begin to increase only at the O2/H2S interface depth. From this same study, activity profiles of dissolved 210Po and 210 Pb reveal that Po and Pb are sequestered efficiently by particulates in surface waters but are remobilized during the reductive dissolution of the carrier phase oxyhydroxides in the aerobic manganese reduction (AMR) zone. Both 210Po and 210Pb are highly enriched at the O2/H2S interface which is also the location of a distinctive microbial layer (primarily anoxygenic phototrophs). The coincident peaks in 210Po, 210Pb and microbial biomass suggest a strong biological influence on the behavior of these radionuclides. There is a strong covariance between the vertical distribution of Mn and Pb, indicating that their redox cycling is closely coupled and is likely microbially mediated (Swarzenski et al., 1999a). Vertical distributions of dissolved Ba and total (dissolved + particulate) Pu, Am and Th in Framvaren Fjord all show increased concentrations with depth (Falkner et al., 1993; Roos et al., 1993). Ba cycling was dominated by its uptake into particulate matter associated with productivity in surface waters, followed by its regeneration at depth or in the sediments. Microbiological activity near the redox interface likely promotes the breakdown of settling particulate matter and the release of barite just above the O2/H2S interface (Falkner et al., 1993). Complex formation with dissolved organic carbon (DOC) is believed to be the main cause for the observed behavior of Pu, Am and Th (Roos et al., 1993). The distributions of these elements were not examined within the regions near the O2/H2S interface and the associated microbial layer. High spatial resolution sampling (especially near the chemocline) in Framvaren Fjord revealed complex and dynamic biogeochemical features that had not previously been observed. McKee and Todd (1993) observed that dissolved 238U concentrations in Framvaren Fjord were approximately 60% lower at the bacterial maximum layer (24 m), situated a few metres below the oxic– anoxic (i.e. oxygen–hydrogen sulfide) interface (18 m), than concentrations above and below this depth. Particulate 238U concentration profiles were a mirror image of the dissolved profiles. During the sampling period (June 1989), removal of U occurred well below the depth at which Fe and Mn oxyhydroxides are precipitated and under conditions inconsistent with abiological reduction of soluble U(VI) to particle-reactive U(IV). These observations suggest that the microbial population in the anoxic waters near the O2/H2S interface in the fjord exerts an effective control on the aquatic biogeochemistry of U in this environment. In contrast, during August 1995, a pronounced peak in both particulate (>0.2 μm) and dissolved (<0.2 μm) uranium was observed at the

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O2/H2S interface (Swarzenski et al., 1999b). Such concentration maxima at the redox boundary is also observed for DOC, Sr and Ba. The authors hypothesize that the source of elevated U at the redox boundary must be due to microbial uptake and subsequent release processes. Uranium oxidation state determinations in waters from 1, 22 and 30 m depth reveal that reduced U(IV) is not present in significant abundance, and that the chemical and/or biological reduction of hexavalent uranium is largely inhibited. These results suggest that U, DOC, Sr, Ba, Fe(II), and Mn(II) are greatly modified by direct and indirect microbial transformation reactions which are most concentrated across the redox transition zone in Framvaren Fjord. Runoff is the main source for the particulate matter to Framvaren Fjord and supplies the surface water of Framvaren with silica, aluminum, manganese, iron, copper, zinc, cadmium and lead, primarily resulting from precipitation of acid rain containing some sea salts, and the dissolution of Farsundite (granitic rock) within the catchment area (Dyrssen, 1997). Geochemical studies of the trace metal concentrations in SPM and sediment trap material from Framvaren, in 1989 and 1993, indicate that extremely high concentrations of zinc, copper, lead (and sometimes cadmium) occur in the SPM collected in the anoxic water layer. The highest concentrations of Zn occur just below the redoxcline at 22 m water depth (in 1993), and copper, lead and cadmium have maximum concentrations between 30 and 80 m depth, where the amount of total SPM is at a minimum. On a mass per volume basis, the maximum concentrations of Zn occur just below the redoxcline (22 m depth), whereas the maximum concentrations of Cd, Cu and Fe occur approximately 1 m above the interface (Skei et al., 1996). The sharp concentration peak of the metals Cu, Cd and Zn at the redoxcline of the Framvaren Fjord, is described as an effect of chelation/ complexation with organic ligands. The most dominating ligand is 2-mercaptobenzothiazole (MBT), which binds with a 2:1 ligand to metal stoichiometric ratio. MBT has the highest concentrations at the vicinity of the redoxcline where its production is likely due to sulfide oxidizing bacteria. Nickel does not exhibit the same distribution as the other three metals due to lower preference for coordination with S and N donor atoms and a much lower rate for loss of water than Cu, Cd and Zn (Hallberg & Larsson, 1999). 3.3.4

Fe–S systematics

When the organic matter produced in surface waters from photosynthesis reaches anoxic marine environments, regeneration is largely coupled with sulfate reduction. Hydrogen sulfide from microbial sulfate reduction reacts with Fe(II) through various mechanisms to produce several iron sulfides, of which pyrite (FeS2) is the most stable in marine sediments. It is possible that the presence of sedimentary pyrite in the geologic record can be used as an indicator of bottom water conditions. However, it has been argued that pyrite

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can be produced in the anoxic water column (Skei, 1988b) as well as formed in the underlying sediments. Until this question of formation is better resolved, the utility of pyrite as a paleoceanographic tool is also in question. It is well documented that sedimentary pyrite formation proceeds through an iron monosulfide (FeS) precursor. Laboratory studies have traditionally indicated that intermediate S species such as elemental sulfur or polysulfides are responsible for the transformation of FeS to FeS2. Recent experimental work however, has suggested that H2S may also be responsible for the transformation. Sediments of Effingham Inlet (a fjord on Vancouver Island) are characterized by high concentrations of FeS, but rapid deposition/burial decouples FeS formation from the formation of intermediate S species and, as a result, H2S is responsible for the transformation of FeS to FeS2 (Hurtgen et al., 1999). The authors hypothesize that on decadal timescales, H2S (or HS−) is responsible for transformations of FeS into FeS2 in natural systems where intermediate S species are isolated from FeS production. The bottom sediments of Framvaren Fjord are enriched in highly reactive iron (Raiswell & Canfield, 1998), and iron monosulfides and pyrite are unusually enriched in 34S (Saelen et al., 1993). Most of these sulfides probably form in bottom waters, where reactive detrital iron phases encounter isotopically heavy dissolved sulfide. Enrichment in 34S in Framvaren is a result of intense bacterial reduction of a sulfate reservoir whose exchange with the open sea is limited by shallow sills and a low-salinity surface-water layer. The Framvaren data indicate that isotopically heavy pyrite forms in anoxic bottom waters. Framvaren sediments contain a small lithogenous fraction and a large biogenous, organic C-rich fraction, which decays by sulfate reduction in an iron-rich water column to form pyrite-rich sediment. The oxidation of H2S and sulfite and the formation of thiosulfate vary according to the concentration of Fe(II) and Mn(II) in Framvaren Fjord (Millero, 1991). The deep waters contained variable concentration of thiosulfide, undetectable concentrations of sulfite, and concentrations of H2S as high as 6000 μM. Based on the C:N:P ratio observed in organic matter, Wensheng and Millero (1995) estimate that about 30% of sulfide produced by sulfate reduction is removed by processes such as oxidation, formation of FeS2, degassing and incorporation into organic matter. Mandernack and Tebo (1999) utilized a newly developed technique to measure sulfide removal rates at the oxic/anoxic interface in Framvaren Fjord, and found them to be among the highest values reported within the water column of any aquatic habitat. Cutter and Kluckhohn (1999) obtained depth profiles for particulate C, N, S, iron monosulfide (FeS), greigite (Fe3S4), and pyrite in the anoxic water column of Framvaren Fjord. The maximum particulate C and N were found at the oxic/anoxic interface, while the highest concentrations of particulate organic S were in surface waters just above this interface, suggesting production via photosynthesis and relatively rapid regeneration. Pyrite and greigite showed

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concentration maxima just below the oxic/anoxic interface, while FeS was almost non-detectable. Throughout the anoxic water column, iron sulfides comprised 73% of the total particulate sulfur, and filtration experiments indicate that much of the iron sulfide particles were <0.4 μm in size. Changes in hydrographical conditions in Mariager Fjord cause transient non-steady-state events at the chemocline that have a great impact on process rates and the distribution of chemical species there (Zopfi et al., 2001). Maxima of elemental sulfur, thiosulfate and sulfite occurred at the chemocline, but were hardly detectable in the sulfidic deep water. Under non-steady-state conditions, where oxic and sulfidic water masses were recently mixed, resulting in an expanded chemocline, the proportion of chemical sulfide oxidation was important; however, under steady state conditions, biological sulfide oxidation (by heterotrophic microorganisms) may account for more than 88% of the total sulfide oxidation. The sulfide oxidation rate determined by incubation experiments was one of the highest reported for stratified basins and about 15 times faster than the rate for chemical oxidation. 3.3.5

Sulfate reduction and methane oxidation

The coincidence of maxima in the methane oxidation rate and the sulfate reduction rate in Saanich Inlet strongly suggests that the methane oxidizing agent was sulfate, either via direct reaction, or coupled indirectly through reactions with other substrates (Devol, 1983). A methane–sulfate coupled reaction diffusion model was developed to describe the inverse relationship commonly observed between methane and sulfate concentrations in the pore waters of anoxic marine sediments. When fit to data from Saanich Inlet (B.C., Canada) and Skan Bay (Alaska), the model not only reproduces the observed methane and sulfate pore water concentration profiles but also accurately predicts the methane oxidation and sulfate reduction rates. In Saanich Inlet sediments, from 23 to 40% of the downward sulfate flux is consumed in methane oxidation while in Skan Bay this value is only about 12%. Anaerobic methane oxidation is a globally important but poorly understood process. Hoehler et al. (1994) proposed that a consortium of methanogenic bacteria and sulfate reducers is responsible for net oxidation of methane under anoxic conditions, a process called ‘reverse methanogenesis’. Hansen et al. (1998) investigated anaerobic methane oxidation and sulfate reduction in intact marine sediment cores and in headspace-free, undiluted, homogenized, incubation bags. In intact cores, the typical upward concave methane concentration profile indicated methane oxidation in the anoxic part of the sediment. Generally, sulfate reduction rates exceeded methane oxidation rates many-fold. In the sulfate–methane transition zone, sulfate reduction was stimulated compared to rates measured both above and below this zone. Methane oxidation rates determined in incubation bags were equivalent to rates determined in intact sediment cores.

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Methane oxidation rates were proportional to the concentrations of methane, and also increased with increasing methane concentrations in the absence of sulfate or the presence of molybdate. When sulfate was added to sulfatedepleted incubation bags, methane oxidation rates decreased immediately to less than half the rate measured prior to the addition, while sulfate reduction was stimulated. When molybdate (a specific inhibitor of sulfate-reducing bacteria) was added to a sulfate-free incubation bag, methane oxidation responded after a lag period of approximately 3 days by uncoupling methane oxidation rates from methane concentrations. Methane production was not affected. From the outcome of their incubation bag experiments, we conclude that methane is not, as previously proposed, oxidized by sulfate reducers alone. Rates of methane production via both acetate fermentation and CO2 reduction were directly measured with radiotracer techniques in the sulfate-depleted sediments of Saanich and Princess Louisa Inlets (Kuivila et al., 1990). Comparison of measured and modeled rates suggests that these two pathways account for the majority of methane produced below the sulfate reduction zone in the sediments of both the basins. Prior aerobic degradation of the organic matter has little influence on the pathways of methane production. 3.3.6

Elemental cycling in sediments underlying anoxic waters

Sediments underlying anoxic bottom waters contain particles that have run the gauntlet of water column diagenesis or were formed in the water column. Measurements of bulk metal concentrations in Saanich Inlet sediments (Francois, 1988) revealed that Ba, Ni, V, Cr, Zn, Pb, Cu and Mo were enriched over the possible contributions from lithogenic sources. Some enrichments (Ba, Cr and Zn) were associated with opaline silica or planktonic materials, indicating an authigenic source in the water column. Other enrichments (Ni, V and Mo) were due to reactions occurring at the sediment–water interface. Uranium concentrations and 234U/238U ratios of particulate matter collected by sediment traps and U concentrations in sediment pore waters (Anderson et al., 1989) indicate that dissolved U is incorporated into particles in surface waters, possibly in association with autochthonous organic matter and also diffuses into the sediments of Saanich Inlet where it is precipitated. Research in anoxic environments has suggested that some minor and trace elements are precipitated where free dissolved sulfide is present without undergoing a valency change, whereas others undergo a change in valency and are either more efficiently adsorbed onto solid surfaces under oxic or anoxic conditions or are precipitated under anoxic conditions. Therefore, the enrichment of many trace elements relative to their crustal abundances indicates only that the sediments accumulated under anoxic conditions, although not necessarily under anoxic bottom waters. However, Calvert and Pedersen (1993) demonstrated that I and Mn enrichments are reliable indicators of bottom

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water oxygenation. Other trace element concentrations or ratios in bottom sediments have been used successfully, in a retrospective manner, to provide insight about the oxic conditions of overlying waters during deposition. The trace elements Re and Mo are strongly enriched in reducing sediments, due to their high ratio of [metal](sw)/[metal](crust), but no enrichment is observed in sediments overlain by oxic waters (Crusius et al., 1996). Therefore, these two trace elements can also be used to infer historical redox conditions even where a substantial lithogenic component can obscure authigenic enrichments of other metals. Russell and Morford (2001) used well-defined peaks in Mo/Al, Cd/Al and Re/Al and a broad peak in U/Al in the lower portion of a debris flow layer in Saanich Inlet to determine that this massive layer was deposited under anoxic bottom waters. Morford et al. (2001) used the sediment concentrations of redox-sensitive metals (Mn, V, Mo, U, Cd and Re) in Saanich Inlet, to determine changes in redox conditions during the past 12.5 kyr. 3.3.7

Preservation of organic matter

A comparison of organic matter (OM) components in cores from Dabob Bay (oxic bottom waters) and Saanich Inlet (anoxic bottom waters) indicates that O2 availability ultimately has little or no independent effect on OM preservation in these environments (Cowie & Hedges, 1992). Comparative analyses of organic compounds in sediment traps and bottom sediments from Saanich Inlet indicates that the anoxic benthic interface is an important site of diagenesis, and that selective removal takes place at both compound-class and molecular levels (Cowie et al., 1992). Preferential loss of marine organic material is indicated by the calculated delta-C-13 value and biochemical composition of the substrate. A better understanding of organic matter preservation in anoxic fjords could lead to a more effective use of these environments as records of historical changes in paleoproductivity or watershed land use. Filippelli (2001) demonstrated that Corg/Porg ratios are high and increase with depth in the Saanich Inlet, but this effect is due largely to a remobilization of P from an organic matter sink to an authigenic sink, and does not reflect changes in paleoproductivities or organic matter inputs to the fjord. A 130-year record of organic carbon content and stable carbon isotope composition in Saanich Inlet sediments (Tunnicliffe, 2000) indicates that organic carbon content increased from 2.5% in the 1860s to 3.3% in the 1990s, whereas carbon isotopes become depleted by 1.2 parts per thousand towards the late 1990s. These findings are consistent with known changes in logging rates within the watershed on lands from which most of the detrital input to the fjord originates. On a similar core, Villanueva and Hastings (2000) examined the early diagenetic processes of chlorophyll a alteration during the previous 157 years. Excellent preservation of pigments, as indicated by high total pigment concentrations, and the presence

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of labile compounds such as chlorophyllide a, indicating that no detectable alteration of the pigment composition has occurred during this time. This suggests that these pigments and biomarkers can be excellent indicators of paleoproductivities in anoxic fjords like Saanich Inlet. Framvaren has a well documented history of environmental change from fjord to lake and then back to a fjord. Sediment cores from both oxic and anoxic past environments contain organic carbon contents of 6–18% and reveal complex distributions of lipid compounds, the dominant classes being n-alkanes, n-alcohols, sterols and long-chain alkenones (Ficken & Farrimond, 1995). The alkenones, which are predominantly produced by marine prymnesiophyte algae, are most abundant in the fjord facies of both cores, but are also detectable in the lacustrine sediments, albeit in much lower concentrations. The dramatic increase in abundance of these compounds is interpreted to record the change in environment from lake to fjord. These studies present great promise that bottom sediments in anoxic fjords may hold important keys to unlocking the past, particularly with regards to global climate change history. References Anderson, R.F., Lehuray, A.P., Fleisher, M.Q. and Murray, J.W. (1989) Uranium deposition in Saanich Inlet sediments, Vancouver Island. Geochimica et Cosmochimica Acta, 53, 2205–2213. Anschutz, P., Sundby, B., Lefrançois, L., Luther III, G.W. and Mucci, A. (2000) Interactions between metal oxides and species of nitrogen and iodine in bioturbated marine sediments. Geochimica et Cosmochimica Acta, 64, 2751–2763. Bothner, M.H., Jahnke, R.A., Peterson, M.L. and Carpenter, R. (1980) Rate of mercury loss from contaminated estuarine sediments. Geochimica et Cosmochimica Acta, 44, 273–285. Bullister, J.L. and Bing-Sun, L. (1995) Chlorofluorocarbon-11 removal in anoxic marine waters. Geophysical Research Letters, 22, 1893–1896. Calvert, S.E. and Pedersen, T.F. (1993) Geochemistry of recent oxic and anoxic marine sediments – implications for the geological record. Marine Geology, 113, 67–88. Cowie, G.L. and Hedges, J.I. (1992) The Role of Anoxia in Organic-Matter Preservation in Coastal Sediments – Relative Stabilities of the Major Biochemicals Under Oxic and Anoxic Depositional Conditions. Organic Geochemistry, 19, 229–234. Cowie, G.L., Hedges, J.I. and Calvert, S.E. (1992) Sources and Relative Reactivities of Amino-Acids, Neutral Sugars, and Lignin in an Intermittently Anoxic Marine Environment. Geochimica et Cosmochimica Acta, 56, 1963–1978. Crusius, J., Calvert, S., Pedersen, T. and Sage, D. (1996) Rhenium and molybdenum enrichments in sediments as indicators of oxic, suboxic and sulfidic conditions of deposition. Earth and Planetary Science Letters, 145, 65–78. Cutter, G.A. and Kluckhohn, R.S. (1999) The cycling of particulate carbon, nitrogen, sulfur, and sulfur species (iron monosulfide, greigite, pyrite, and organic sulfur) in the water columns of Framvaren Fjord and the Black Sea. Marine Chemistry, 67, 149–160. Deflandre, B., Mucci, A., Gagné, J.-P., Guignard, C. and Sundby, B. (2002) Early diagenetic processes in coastal marine sediments disturbed by a catastrophic sedimentation event. Geochimica et Cosmochimica Acta, 66(16), 2547–2558. Devol, A.H. (1983) Methane oxidation rates in the anaerobic sediments of Saanich Inlet. Limnology and Oceanography, 28, 738–742.

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Dyrssen, D.W. (1997) The source of silica and trace metals in particulate matter in Framvaren fjord. Aquatic Geochemistry, 3, 329–343. Dyrssen, D.W. (1999) Framvaren and the Black Sea – Similarities and differences. Aquatic Geochemistry, 5, 59–73. Dyrssen, D.W., Skei, J.M., Ostlund, H.G., Hall, P.O.J., Haraldsson, C. and Chierici, M. (1996) Time dependence of organic matter decay and mixing processes in Framvaren, a permanently anoxic fjord in south Norway. Aquatic Geochemistry, 2, 111–129. Edenborn, H.M., Mucci, A., Silverberg, N. and Sundby, B. (1987) Sulfate reduction in deep coastal marine sediments. Marine Chemistry, 21, 329–345. Falkner, K.K., Klinkhammer, G.P., Bowers, T.S., Todd, J.F., Lewis, B.L., Landing, W.M. and Edmond, J.M. (1993) The behavior of barium in anoxic marine waters. Geochimica et Cosmochimica Acta, 57, 537–554. Ficken, K.J. and Farrimond, P. (1995) Sedimentary lipid geochemistry of Framvaren: impacts of a changing environment. Marine Chemistry, 51, 31–43. Filippelli, G.M. (2001) Carbon and phosphorus cycling in anoxic sediments of the Saanich Inlet, British Columbia. Marine Geology, 174, 307–321. Francois, R. (1988) A study on the regulation of the concentrations of some trace metals (Rb, Sr, Zn, Pb, Cu, V, Cr, Ni, Mn and Mo) in Saanich Inlet sediments, British Columbia, Canada. Marine Geology, 83, 285–308. Freeland, H.J., Farmer, D.M. and Levings, C.D. (eds) (1979) Fjord Oceanography, NATO Conference Series, Series IV: Marine Sciences, vol 4, 715 pp. Gagnon, C., Mucci, A. and Pelletier, E. (1995) Anomalous accumulation of acid-volatile sulphides (AVS) in a coastal marine sediment, Saguenay Fjord, Canada. Geochimica et Cosmochimica Acta, 59, 2663–2675. Gagnon, C., Mucci, A. and Pelletier, E. (1996a) Vertical distribution of dissolved sulphur species in coastal marine sediments. Marine Chemistry, 52, 195–210. Gagnon, C., Mucci, A., Pelletier, E. and Fitzgerald, W.F. (1996b) The diagenetic behaviour of methylmercury in a contaminated organic-rich coastal marine sediment. Limnology and Oceanography, 41, 428–434. Gagnon, C., Pelletier, E. and Mucci, A. (1997) The behaviour of anthropogenic mercury in coastal marine sediments. Marine Chemistry, 59, 159–176. German, C.R. and Elderfield, H. (1989) Rare earth elements in Saanich Inlet, British Columbia, a seasonally anoxic basin. Geochimica et Cosmochimica Acta, 53, 2561–2571. Gobeil, C. and Cossa, D. (1993) Mercury in sediments and sediment pore water in the Laurentian Trough. Canadian Journal of Fisheries and Aquatic Sciences, 50, 1794–1800. Gregory, J.W. (1913) The Nature and Origin of Fjords, John Murray, London. Grill, E.V. (1982) Kinetic and thermodynamic factors controlling manganese concentrations in oceanic waters. Geochimica et Cosmochimica Acta, 46, 2435–2446. Hallberg, R.O. and Larsson, C. (1999) Biochelates as a cause of metal cycling across the redoxcline. Aquatic Geochemistry, 5, 269–280. Hansen, L.B., Finster, K., Fossing, H. and Iversen, N. (1998) Anaerobic methane oxidation in sulfate depleted sediments: effects of sulfate and molybdate additions. Aquatic Microbial Ecology, 14, 195–204. Hoehler, T.M., Alperin, M.J., Albert, D.B. and Martens, C.S. (1994) Field and laboratory studies of methane oxidation in an anoxic marine sediment: evidence for a methanogen-sulfate reducer consortium. Global Biogeochemical Cycles, 8, 451–463. Holtedahl, H. (1967) Notes on the formation of fjords and fjord – valleys. Geografiska Annaler, 49, 188–203. Hulth, S., Aller, R.C. and Gilbert, F. (1999) Coupled anoxic nitrification/manganese reduction in marine sediments. Geochimica et Cosmochimica Acta, 63(1), 49–66. Hurtgen, M.T., Lyons, T.W., Ingall, E.D. and Cruse, A.M. (1999) Anomalous enrichments of iron monosulfide in euxinic marine sediments and the role of H2S in iron sulfide transformations: Examples from Effingham Inlet, Orca Basin, and the Black Sea. American Journal of Science, 299, 556–588.

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Kuivila, K.M., Murray, J.W. and Devol, A.H. (1990) Methane production in the sulfate-depleted sediments of two marine basins. Geochimica et Cosmochimica Acta, 54, 403–411. Lee, B.S., Bullister, J.L. and Whitney, F.A. (1999) Chlorofluorocarbon CFC-11 and carbon tetrachloride removal in Saanich Inlet, an intermittently anoxic basin. Marine Chemistry, 66, 171–185. Lidstrom, M.E. (1983) Methane consumption in Framvaren, an anoxic marine fjord. Limnology and Oceanography, 28, 1247–1251. Loring, D.H. (1975) Mercury in the sediments of the Gulf of St. Lawrence. Canadian Journal of Earth Sciences, 12, 1219–1237. Loring, D.H. and Bewers, J.M. (1978) Geochemical mass balances for mercury in a Canadian fjord. Chemical Geology, 22, 309–330. Louchouarn, P. and Lucotte, M. (1998) A historical reconstruction of organic and inorganic contamination events in the Saguenay Fjord/St. Lawrence system from preindustrial times to the present. The Science of the Total Environment, 2213, 139–150. Louchouarn, P., Lucotte, M., Canuel, R., Gagné, J.-P. and Richard, L.F. (1997) Sources and early diagenesis of lignin and bulk organic matter in the sediments of the lower St. Lawrence Estuary and the Saguenay Fjord. Marine Chemistry, 58, 3–26. Mandernack, K.W. and Tebo, B.M. (1999) In situ sulfide removal and CO2 fixation rates at deep-sea hydrothermal vents and the oxic/anoxic interface in Framvaren Fjord, Norway. Marine Chemistry, 66, 201–213. McKee, B.A. and Todd, J.F. (1993) Uranium Behavior in a Permanently Anoxic Fjord – Microbial Control. Limnology and Oceanography, 38, 408–414. Millero, F.J. (1991) The oxidation of H2S in Framvaren Fjord. Limnology and Oceanography, 36, 1007–1014. Morford, J.L., Russell, A.D. and Emerson, S. (2001) Trace metal evidence for changes in the redox environment associated with the transition from terrigenous clay to diatomaceous sediment, Saanich Inlet, BC. Marine Geology, 174, 355–369. Mucci, A. and Edenborn, H.M. (1992) Influence of an organic-poor landslide deposit on the early diagenesis of iron and manganese in a coastal marine sediment. Geochimica et Cosmochimica Acta, 56, 3909–3921. Mucci, A., Richard, L.F., Lucotte, M. and Guignard, C. (2000a) The differential geochemical behavior of arsenic and phosphorus in the water column and sediments of the Saguenay Fjord Estuary, Canada. Aquatic Geochemistry, 6, 293–324. Mucci, A., Sundby, B., Gehlen, M., Arakaki, T. and Silverberg, N. (2000b) The fate of carbon in continental shelf sediments of Eastern Canada: a case study. Deep-Sea Research II, 47, 733–760. Naes, K., Knutzen, J. and Berglind, L. (1995) Occurrence of PAH in marine organisms and sediments from smelter discharge in Norway. Science of the Total Environment, 163, 93–106. Pelletier, E., Deflandre, B., Nozais, C., Tita, G., Gagné, J.-P., Desrosiers, G. and Mucci, A. (1999) Crue éclair de juillet 1996 dans la région du Saguenay (Québec). 2. Impacts sur les sédiments et le biote de la baie des Ha! Ha! et du fjord du Saguenay. Canadian Journal of Fisheries and Aquatic Sciences, 56, 2136–2147. Raiswell, R. and Canfield, D.E. (1998) Sources of iron for pyrite formation in marine sediments. American Journal of Science, 298, 219–245. Roos, P., Noshkin, V., Ballestra, S., Holm, E., Sanchez, A. and Gastaud, J. (1993) Radioanalytical studies of anthropogenic radionuclides in an anoxic fjord. Science of the Total Environment, 130–131, 1–22. Russell, A.D. and Morford, J.L. (2001) The behavior of redox-sensitive metals across a laminated-massive-laminated transition in Saanich Inlet, British Columbia. Marine Geology, 174, 341–354. Saelen, G., Bottrell, S.H., Raiswell, R., Talbot, M.R. and Skei, J.M. (1993) Heavy sedimentary sulfur isotopes as indicators of super-anoxic bottom-water conditions. Geology, 21, 1091–1094.

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Sanchez, A.L., Gastaud, J., Holm, E. and Roos, P. (1994) Distribution of plutonium and its oxidation states in Framvaren and Hellvik fjords, Norway. Journal of Environmental Radioactivity, 22, 205–217. Saulnier, I. and Mucci, A. (2000) Trace metal remobilization following the resuspension of estuarine sediments: Saguenay Fjord, Quebec. Applied Geochemistry, 15, 191–210. Schafer, C.T., Smith, J.N. and Loring, D.H. (1980) Recent sedimentation events at the head of the Saguenay Fjord, Canada. Environmental Geology, 3, 139–150. Schafer, C.T., Smith, J.N. and Seibert, G. (1983) Significance of natural and anthropogenic sediment inputs to the Saguenay Fjord, Quebec. Sedimentary Geology, 36, 177–194. Shapiro, S.D., Schlosser, P., Smethie, W.M. and Stute, M. (1997) The use of H-3 and tritiogenic He-3 to determine CFC degradation and vertical mixing rates in Framvaren Fjord, Norway. Marine Chemistry, 59, 141–157. Skei, J.M. (1982) Permanently anoxic, marine basins – exchange of substances across boundaries, in Environmental Biogeochemistry, Ecological Bulletin, vol. 35 (ed. R. Hallberg), pp. 419–424. Skei, J.M. (1983) Why sedimentologists are interested in fjords. Sedimentary Geology, 36, 75–80. Skei, J.M. (1988a) Framvaren – Environmental setting. Marine Chemistry, 23, 209–218. Skei, J.M. (1988b) Formation of framboidal iron sulfide in the water of a permanent anoxic fjord – Framvaren. South Norway. Marine Chemistry, 23, 345–353. Skei, J.M., Bakke, T. and Molvaer, J. (2000) The Norwegian Coast, in Seas at the Millennium. An Environmental Evaluation (ed. C. Sheppard), Elsevier Science Ltd, pp. 17–30. Skei, J.M., Larsson, P., Rosenberg, R., Jonsson, P., Olsson, M. and Broman, D. (2000). Eutrophication and Contaminants in Aquatic Ecosystems. Ambio, 29, 184–194. Skei, J.M., Loring, D.H. and Rantala, R.T.T. (1996) Trace metals in suspended particulate matter and in sediment trap material from a permanently anoxic fjord – Framvaren, south Norway. Aquatic Geochemistry, 2, 131–147. Smith, J.N. (1988) Pollution history and paleoclimatic signals in sediments of the Saguenay Fjord. In Chemical Oceanography in the Gulf of St. Lawrence, vol. 220 (ed. P.M. Strain), pp. 123–138. Smith, J.N. and Loring, D.H. (1981) Geochronology for mercury pollution in the sediments of the Saguenay Fjord. Environmental Science and Technology, 15, 944–951. Smith, J.N. and Schafer, C.T. (1987) A 20th century record of climatologically-modulated sediment accumulation rates in a Canadian fjord. Quaternary Research, 27, 232–247. Smith, J.N. and Walton, A. (1980) Sediment accumulation rates and geochronologies measured in the Saguenay Fjord using the Pb-210 dating method. Geochimica et Cosmochimica Acta, 44, 225–240. Sternbeck, J., Skei, J., Verta, M. and Østlund, P. (1999) Mobilisation of sedimentary trace metals following improved oxygen conditions – an assessment of the impact of a lowered primary productivity on trace metal cycling in the marine environment. TemaNord, 594, 65 pp. St-Onge, G. and Hillaire-Marcel, C. (2001) Isotopic constraints of sedimentary inputs and organic carbon burial rates in the Saguenay Fjord, Quebec. Marine Geology, 176, 1–22. Sundby, B., Gobeil, C., Silverberg, N. and Mucci, N. (1992) The phosphorus cycle in coastal marine sediments. Limnology and Oceanography, 37, 1129–1145. Swarzenski, P.W., McKee, B.A., Skei, J.M. and Todd, J.F. (1999a) Uranium biogeochemistry across the redox transition zone of a permanently stratified fjord: Framvaren, Norway. Marine Chemistry, 67, 181–198. Swarzenski, P.W., McKee, B.A., Sorensen, K. and Todd, J.F. (1999b) Pb-210 and Po-210, manganese and iron cycling across the O2/H2S interface of a permanently anoxic fjord: Framvaren, Norway. Marine Chemistry, 67, 199–217. Syvitski, J.P.M. (1984) Sedimentology of Arctic Fjords Experiment: HU83-028 data report, vol. 2, Canadian Data Report of Hydrography and Ocean Sciences, 28, 1130 pp. Syvitski, J.P.M. and Schafer, C.T. (1996) Evidence for an earthquake-triggered basin collapse in Saguenay Fjord, Canada. Sedimentary Geology, 104, 127–153.

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4

The Eastern Mediterranean Michael D. Krom, Steven Groom and Tamar Zohary

4.1

Introduction

The Mediterranean in general and the Eastern Mediterranean in particular is a unique ocean. It is situated immediately south of industrialised Europe with a population of approximately 120 million people living close to its shoreline (Turley, 1999) and is subjected to an annual invasion of a further 120–200 million tourists. Despite the natural inputs of nutrients to the system and the considerable additional anthropogenic pressures which this population subjects the basin to, it is a marine desert. It is as much of a desert as the Sahara which is located immediately to the south of the basin. Many studies have shown that the Eastern Mediterranean is ultra-oligotrophic and probably the most oligotrophic sea known. The deep blue colour of the pelagic water is an expression of this extreme desert nature. The world record Secchi depth, of 53 m, has been reported from the southern Levantine Basin (Berman et al., 1985). Other lines of evidence which demonstrate the ultra-oligotrophic nature of the system include the low depth-integrated chlorophyll values (Yacobi et al., 1995), the high contribution of pico- and nano-phytoplankton to this total chlorophyll (Li et al., 1993; Vidussi et al., 2001), and the deeper deep chlorophyll maximum (DCM) ranging seasonally between 90–110 m in summer-fall and 130–140 m in winter-spring. These and related issues will be discussed in more detail below. The concentration of dissolved nutrients in the deep waters of the Eastern Mediterranean is much lower than those in other oceans of the world, and when these are mixed into the surface waters they support very low primary productivity. The basic reason for this ultra-oligotrophic status is that the Mediterranean has an anti-estuarine (reverse thermohaline) circulation in which nutrientdepleted surface waters flow into the western basin at the Straits of Gibraltar and then on into the eastern basin at the Straits of Sicily. The deeper counter current consists of Levantine Intermediate Water (LIW) which contains a significant amount of dissolved nutrients. In addition to this net export of nutrients, the Eastern Mediterranean is the largest body of water in the world in which the primary productivity is phosphorus limited (Krom et al., 1991). The deep waters of the eastern basin have a nitrate/phosphate ratio of 25–28:1 compared to 22:1 in the Western Mediterranean and 15–17:1 found generally in the world’s oceans. These

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unique features combine to create an unusual combination of biogeochemical features, which is the subject of this chapter.

4.2

History of the Mediterranean basin

The Mediterranean Sea is an ocean basin that is gradually closing as the African Plate moves north and collides with Europe. It is the remains of the far larger and more extensive Tethys ocean. In the comparatively recent geological past (Late Miocene, 7.5–5 million years ago), there were periods when the basin was completely dried out. This has resulted in the extensive deposits of salt that underlie much of the sea-bed (and surrounding areas) of the present-day Mediterranean. It is the source of the salt found in the several small saline pools (e.g. Bannock Basin) which are found within the Eastern Mediterranean basin. During the Holocene and Pleistocene, circulation in the Eastern Mediterranean became so restricted at times that the deeper waters became anoxic as in the Black Sea today. During these periods of restricted circulation, the primary productivity of the basin increased (Calvert et al., 1992) and a series of organic-rich sediments called sapropels were deposited (de Lange et al., 1999). These sapropel deposits are found interlayered with the organic-poor calcareous silts and muds. The calcareous silts and muds are similar to those that are deposited in the basin today and were presumably deposited under similar oceanographic and biogeochemical conditions. The precise paleo-oceanographic conditions that gave rise to sapropel deposits are still actively debated. It is generally agreed that during the periods when sapropels were formed, the climate around the basin was less arid than it is today resulting in less net evaporation (Bar-Matthews et al., 1999). Sapropel deposition also coincided with periods when there were increased rainfall in central and eastern Africa and far higher river flow in the Nile (Rossignol-Strick et al., 1982). Together these conditions changed the circulation pattern in the basin. Whether the circulation remained anti-estuarine (Myers et al., 1998) or whether it is a requirement of sapropel deposition that the circulation must have become estuarine, at least for part of the basin (Struck et al., 2001) is still discussed. Such a change in circulation would result in the nutrients being trapped within the basin and available for deposition in the sediment and not exported as is the present situation. In either case, it is clear that natural climatic change has resulted in major changes in the paleo-oceanography of the Eastern Mediterranean. There is good evidence that recent man-induced environmental and climate change is having a noticeable effect on the physical circulation and on the biogeochemistry of this most sensitive of ecosystems that will be discussed at the end of this chapter.

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4.3

Basic description of the bathymetry and physical oceanography of the Eastern Mediterranean

4.3.1

Bathymetry

The Mediterranean is an almost completely enclosed sea with a narrow outlet to the west through the Straits of Gibraltar to the eastern North Atlantic. It can be divided into three major basins (Fig. 4.1). The western basin comprises the region west of Sicily, which includes the Tyrrhenian and Alboran Seas. This area is called the Western Mediterranean. The Eastern Mediterranean comprises a western basin including the Adriatic and Ionian Seas and an eastern basin which includes the Aegean Sea and Levantine Basin and is separated from the Ionian Sea by the Cretan Sea and Cretan Straits. In this review, we will deal mainly with biogeochemical processes in the Aegean Sea and Levantine Basin of the Eastern Mediterranean. 4.3.2

Physical circulation of the Eastern Mediterranean

The Mediterranean is situated at the northern boundary of the desert climatic belt. As a result there is a significant net loss of water from the basin. The Mediterranean in general, and the Eastern Mediterranean in particular, has a large excess of evaporation over precipitation. This results in an anti-estuarine (reverse thermohaline) circulation. Low salinity water flows into the Mediterranean basin from the eastern North Atlantic. This water flows in within the upper water layers at the Straits of Gibraltar. The salinity of this surface water is 36.15. The surface water moves to the west generally following the coast of North Africa. It then flows through the Straits of Sicily. However, the flow of

Adriatic Sea

Black Sea

Western Mediterranean

Straits of Gibraltar

Aegean Sea Straits of Sicily Ionian Sea Levantine Basin

Fig. 4.1 Satellite image of Mediterranean Sea showing the principal basins and straits particularly of the Eastern Mediterranean. The image used is the NASA AVHRR Pathfinder 9-km monthly composite SST image for September 2000. (Data obtained from the NASA Pathfinder project.)

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surface water towards the east is complicated by the many mesoscale eddies and jets that are characteristic of the Mediterranean in general (POEM, 1992). The Straits of Sicily represent a physical barrier which lies at the entrance to the Eastern Mediterranean. It is approximately 500 m deep and represents a significant control on the biogeochemical processes occurring within the eastern basin. There also is an anti-estuarine circulation through these straits because evaporation exceeds precipitation in the eastern basin. A surface and subsurface flow of modified Atlantic waters, low in nutrients, flows eastwards into the eastern basin while LIW with relatively higher nutrients flows westwards out of the basin. This is the fundamental feature that causes the Eastern Mediterranean as a whole to be oligotrophic. The current flow through the straits has been determined recently by current metre measurements (Astraldi et al., 1999). They found that the annual flux of water out of the Strait was 1.1 Sv which is similar to the estimate of total outflow through the Strait of Sicily of 1.2 Sv based on salinity balance (Bethoux, 1981). Astraldi et al. (1999) state that the inflow into the Eastern Mediterranean cannot be measured using current metres, because it is dominated by mesoscale features. They note that in order to maintain the salinity balance of the basin, there must be 4% more water flowing into the Eastern Mediterranean at the Straits of Sicily than are flowing out. It is therefore calculated that the total annual inflow into the Eastern Mediterranean was 1.19 Sv. They also showed that there is a seasonal pattern to the water flow with higher flows in winter than in summer. This has implications for the magnitude and nature of the nutrient fluxes through the straits (see Section 4.10). Figure 4.2 shows a typical vertical temperature and salinity profile from the eastern Levantine Basin in May. Beneath a shallow mixed layer and sharp seasonal thermocline, there is a modified Atlantic water, which has a characteristic low salinity. As the surface water flows to the east the modified Atlantic water becomes progressively more saline. Eventually it is converted into LIW and Levantine Deep Water (LDW). The transformation occurs in rather limited areas within the eastern basin where the oceanographic and climatic conditions combine to cause downward flow of water in winter. As occurs in all parts of the ocean, it results in the downward flow of biogeochemically active chemical substances such as oxygen, carbon dioxide and nutrients from the surface into intermediate and deep water. 4.3.2.1 Formation of LDW Recently there have been changes in the deep water in the Eastern Mediterranean (Lascaratos et al., 1999). Previously LDW was formed entirely in the southern Adriatic Sea with small amounts of additional dense water from the northern Adriatic (Lascaratos et al., 1999). The Adriatic Deep water that exits through the Straits of Ortranto typically has a temperature of 13°C and a salinity of 38.65 (Gacic et al., 1996). This deep water fills the deepest parts of the Ionian

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38.6 0

38.8

39

39.2

13

15

17

19

21

0

Modified Atlantic Water Levantine Intermediate Water

–400 Depth (m)

Depth (m)

–400

–800

–800

Levantine Deep Water –1200

–1200 Aegean Deep Water

–1600

–1600 Salinity

Temperature

Fig. 4.2 A typical T/S profile from recent CYCLOPS cruise to the Southeast Levantine Basin. Data obtained by Dr G. Zodiatis during the CYCLOPS cruise to the Eastern Mediterranean in May 2000.

Sea and the Levantine Basin. The deep waters of these basins used to be very homogenous in temperature and salinity displaying only small horizontal gradients (Fig. 4.2; Schlitzer et al., 1991). Roether and Schlitzer (1991) computed an annual mean formation rate of ~0.3 Sv for LDW formed in the Adriatic. Using freon and tritium measurements, Roether et al. (1996) calculated a mean residence time of LDW in the eastern basin which was 50–80 years. 4.3.2.2 Formation of LIW The intermediate layers of the Eastern Mediterranean (200–500 m) are occupied by saline LIW (Fig. 4.2). LIW exits the Eastern Mediterranean at the Straits of Sicily and is the precursor of the Mediterranean outflow water (MOW) which leaves the Straits of Gibraltar. MOW is warm and saline compared to intermediate water in the North Atlantic. Because of its higher density it plunges to depths of 1000–2000 m in the eastern North Atlantic and contributes to the upwelling of waters at the coast of Spain and Portugal. The precise origin of LIW is still debated. During the summer, the surface layers of the Eastern Mediterranean are occupied by warm salty waters above a sharp seasonal thermocline. Winter cooling increases the density of this surface water until it sinks and forms LIW. This process appears to occur in a number of different locations depending on the particular weather and/or oceanographic features. One location which seems to be important in LIW formation is the Rhodes

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Gyre, which is a permanent cold-core eddy situated to the southwest of the island of Rhodes (Figs 4.3 and 4.7). Lascaratos et al. (1999) suggest that this is the most important single location for LIW formation. Ozsoy et al. (1989) showed that LIW can also form along much of the northern Levantine Sea while the southern Aegean Sea and southern Levantine Basin have also been proposed as sites of LIW formation if the weather and oceanographic conditions are favourable (Wust, 1961; Morcos, 1972; Nittis & Lascaratos, 1998). This inter-annual variability in atmospheric forcing also affects the amount of LIW formed each year. This can vary between 0.6 and 1.3 Sv with a typical climatological average value of ~1.0 Sv. 4.3.3

Recent water mass changes in the Eastern Mediterranean

Starting in 1989, a new source of deep water has been found in the deep Levantine and Ionian Basins. This deep water was formed in the Aegean Sea and then discharged into the Levantine and Ionian Basins via the Straits of Kassos. The new water is denser than the previous LDW formed in the Adriatic. It is warm salty water with a salinity of 39 and a temperature of 14°C. Current speeds in excess of 0.5 m/s were measured in the Straits of Kassos in 1989 (Theocharis et al., 1999), and geostrophic calculations suggest that even higher flow rates probably occurred subsequently. The formation rate of this new deep water is ~1.2 Sv (Roether et al., 1996) which is 4 times the previous rate of formation of deep water formed in the Adriatic. Analysis of the T/S data from the Aegean through the 1980s suggests that initially there was a systematic increase in salinity from 1987–92 and then a decrease in temperature from 1992–95 (Lascaratos et al., 1999). The reason for this new deep water formation is still being debated and is likely to be due to a number of contributory factors. Josey et al. (1997) found that for the period 1988–95, the Aegean was subjected to anomalously high cooling and net evaporation. Theocharis et al. (1999) also present evidence of a significant reduction of precipitation during this period. Josey et al. (1997) also reported that there was a significant decrease in freshwater supply to the entire Eastern Mediterranean during this period as much of the freshwater is intercepted for use in irrigation, industry, etc. The major change in freshwater flow to the Eastern basin was the closure of the Aswan Dam in 1965, which, given the exchange rate of water in the basin should have affected the Aegean Sea at approximately this time. An additional contributory factor seems to be the presence of a blocking anti-cyclonic eddy at the entrance between the Ionian and Levantine Basin. Observations on this new deep water have shown that it contains CFC’s characteristic of water formed recently (Roether et al., 1996). The water has higher dissolved oxygen and lower silicate than the adjacent LDW formed in the Adriatic. Once this deep water flowed into the Levantine Basin, it caused existing deep water to mix upwards into the LIW. Klein et al. (1999) report

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97

a shallowing of the nutricline to a depth of 150 m, which in the Eastern Mediterranean is within the photic zone. They suggested that this should result in a transient increase in primary productivity across the region, which should be followed some time later by a decline as the reservoir of nutrients for upwelling is reduced. No direct observation of such changes has yet been made. This is one of the examples whereby anthropogenically driven climate and environmental change may well have caused significant biogeochemical changes to this basin. 4.3.4

Current patterns

The earliest measurements of physical properties in the water of the Levantine Basin were made from the Norwegian ship RV Thor in 1910. On the basis of these data, the circulation in the upper and intermediate layers of the Levantine Basin was depicted as a simple basin-wide cyclonic gyre. Water flowed in along the North African coast and flowed out along the Turkish and Greek coasts. Data collected subsequently by the American RV Atlantis in 1948, the French RV Calypso in 1956, and various Soviet vessels in the late 1950s through to the early 1970s supported this picture of the large-scale circulation. However during the 1980s, particularly as a result of the work of the POEM (Physical Oceanography of the Eastern Mediterranean) programme it became

Rhodes Gyre Mersah Matru Eddy

Cyprus Eddy

Fig. 4.3 AVHRR 1-km SST image for 24 July 2000 showing mesoscale structure in the Eastern Mediterranean. Raw data obtained from the NOAA Satellite Active Archive and processed at Plymouth Marine Laboratory.

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clear that the current patterns of the Eastern Mediterranean are more complex (POEM, 1992). There are a series of persistent sub-basin-scale gyres interconnected by meandering currents as well as an energetic mesoscale field consisting of warm- and cold-core eddies. Because of the scale of these features, the terms gyre and eddy are often used interchangeably in the Eastern Mediterranean. These mesoscale features can be clearly seen by satellite imagery (Fig. 4.3). Some of these features are very persistent, e.g. the cold-core eddy situated southwest of Rhodes (Rhodes Gyre) and the warm-core eddies south of Cyprus (Cyprus Eddy) and southeast of Crete (Mersah Matru Eddy). These features can have a major effect on the biogeochemical processes occurring within the basin as will be described below. 4.4

4.4.1

Nutrients and chlorophyll distribution across the Eastern Mediterranean General comments

The structure of the water column across the Eastern Mediterranean carries many of the characteristics of a typical tropical structure (Herbland & Voituriez, 1979). However the deep waters of the Eastern Mediterranean have unusually low nutrient concentrations when compared to the deep waters of other oceans. The LDW typically has 6 μmol kg−1 of nitrate, 0.25 μmol kg−1 of phosphate and 10–12 μmol kg−1 of silicate. By comparison, waters of similar depth in the Atlantic Ocean have dissolved nitrate, phosphate and silicate concentrations of 20, 1.8 and 20 μmol kg−1 respectively, and nutrient levels in the deep Indian and Pacific are even higher. 4.4.2

Seasonal distributions

4.4.2.1 Winter In winter, the mixed layer across the Levantine Basin is approximately 150–200m deep (Kress & Herut, 2001) and brings up nutrients accumulated below the DCM during the year. Because the weather in the Mediterranean has relatively warm and sunny days even in winter, the annual phytoplankton bloom occurs as soon as the dissolved nutrients are mixed up from below by deep winter mixing (Fig. 4.4). During this annual bloom, all of the dissolved phosphate is consumed below analytical detection limits while excess nitrate remains at typically ~1 μmol kg−1 (Krom et al., 1992; Kress & Herut, 2001; Struck et al., 2001; Fig. 4.4). A similar pattern has been observed in the Cretan Sea except the mixed layer is somewhat shallower (~100 m) and the residual nitrate in the surface waters after the bloom somewhat higher (~2 μmol kg−1, Tselepides et al., 2000). Silicate levels in the surface layers are typically 2 μmol kg−1 over most of the Levantine Basin and do not change significantly during the year.

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SELevantine: Feb. 1989 0. 00 0

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Chlorophyll a

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SELevantine: May 89

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Chlorophyll

Phosphate

0. 00 0

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SELevantine: Nov. 89

SELevantine: Nov. 89 0

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– 400 – 600

–800

– 800

–1000

– 1000

Phosphate

Chlorophyll

Nitrate

Temp.

Fig. 4.4 Seasonal nutrient distribution in Southeast Levantine Basin (after Krom et al., 1992).

This is because in general, in the open waters the dominant organisms involved in the phytoplankton bloom are nano- and pico-plankton which do not have a siliceous skeleton. There are few siliceous phytoplankton within the winter phytoplankton bloom and even less during the spring–summer period which follows (see Section 4.4.2.2).

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4.4.2.2 Spring into summer As soon as the permanent seasonal thermocline develops and deep mixing ceases, a DCM forms. Nitrate levels in the surface waters are progressively reduced until by summer (May) all of the nitrate has been consumed (Fig. 4.4; Kress & Herut, 2001). From May until deep mixing occurs again in winter, the inorganic nutrients (nitrate, ammonia and phosphate) in the surface waters are below analytical detection limits and only accumulate below the DCM (Fig. 4.4). Krom et al. (in preparation) speculate that this consumption of nitrate, while phosphate levels remain below detection limits, may be due in part to the consumption of dissolved organic phosphorus by phytoplankton and bacteria. The CYCLOPS team have found high levels of dissolved organic phosphorus (DOP) (both refractory and UV-labile DOP) in surface waters of the Levantine Basin (Carbo et al., 2002). The amounts found (50–400 nmol kg−1) are quite sufficient to consume all of the excess nitrate remaining in the water column after the winter phytoplankton bloom. Recent work in the North Central Pacific and North Atlantic (Sargasso Sea) have shown that dissolved organic nitrogen (DON) and DOP can be important in oligotrophic waters in controlling both the nature and amount of primary productivity occurring (Wu et al., 2000). Alternatively, the consumption of nitrate could be brought about by efficient recycling of phosphorus in the photic zone while nitrogen-rich organic matter gets exported to the lower layers. 4.4.3

Nutrient distribution below the nutricline

Across the Levantine Basin, the nutrients increase with depth to a maximum of 6 μmol kg−1 of nitrate, 0.25 μmol kg−1 of phosphate and 10–12 μmol kg−1 of silicate at 500–600 m towards the bottom of the LIW (Kress & Herut, 2001). Below that depth, nutrients are generally constant within the LDW or sometimes decrease by a small amount (Figs 4.4 and 4.8). At the same time, dissolved oxygen decreases from slightly oversaturated levels in the DCM to levels of ~70% saturation at depth. There is no oxygen minimum and little or no nutrient maximum in the open Eastern Mediterranean, and the bottom waters never approach oxygen levels where denitrification or any other anaerobic process occurs. Indeed, it is characteristic of the sediments of the Eastern Mediterranean that at present they are entirely aerobic calcareous muds with the exception of the northern Adriatic and certain other limited nearshore areas. The redox boundary, such as it is, is generally 10–20 cm within the sediment and is a relic of the period of sapropel times, more than 5000 years ago.

4.5

Total chlorophyll distribution and characteristics

The DCM, which is characteristic of the non-winter situation in the Eastern Mediterranean, typically develops at a depth of 100–140 m (Fig. 4.5; Yacobi

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Chlorophyll a, ng l–1 0

40 100 150 200

80 250

200

120

200 150 100

100

100 160

50

28.5

29.0

29.5

30.0

30.5

31.0

31.5

32.0

32.5

Fig. 4.5 Chlorophyll concentrations on an east–west (33°30′N) transect across the southern Levantine Basin. Stations were located at intervals of 0.5. At each station samples were taken from the following depths: 0, 25, 50, 75, 90, 100, 110, 130, 150, 170 and 200 m (figure from Yacobi et al., 1995).

et al., 1995). Different studies of the nature of the DCM have reported that it coincides with the maximum in total fluorescence but not with the maximum abundance of cells nor with the presumed maximum in the carbon biomass of the ultra-phytoplankton (Kimor et al., 1987; Li et al., 1993; Yacobi et al., 1995; Zohary et al., 1998; Psarra et al., 2000). Those studies also show that the cell concentrations above the DCM are usually much more uniform with depth than chlorophyll concentrations. This discrepancy is explained by light adaptation of the algal cells which lead to a higher cellular content of chlorophyll at the deeper depths than in the overlying water, permitting cells to photosynthesise even at 100 m depth and at 1% of the surface irradiance. Interestingly, various studies (Yilmaz et al., 1993; Yacobi et al., 1995; Yilmaz & Tugrul, 1998) have pointed out that the DCM tends to be limited to a narrow range of water densities between 28 and 28.6. An important characteristic of the phytoplankton of the Eastern Mediterranean is their small size. Most of the chlorophyll (>60 to >98%, depending on the study) is associated with particles smaller than 10 μm, or even 3 μm (Li et al., 1993; Zohary et al., 1998; Psarra et al., 2000; Vidussi et al., 2001). This abundance in the smaller size range is typical of oligotrophic seas and is thought to be due to the greater ability of small phytoplankton to harvest the very low levels of nutrients available in such systems due to their higher surface area to volume ratio.

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Light penetration

The deep blue colour and the exceptionally high transparency of the Eastern Mediterranean means that light penetrates deeper into these waters, making the euphotic zone (i.e. the illuminated upper layer) thicker. This allows photosynthetic organisms to thrive at greater depths than in other less transparent oceans. Usually, marine phototrophs are limited by light to the depth receiving 0.1–1.0% of surface irradiance. In the Eastern Mediterranean the euphotic zone generally extends to greater than 150 m and often to 200 m. 4.5.2

Species composition

The phytoplankton of the pelagic waters of the Eastern Mediterranean are typically composed of three major functional groups which are also distinct taxonomically. These are the prochlorophytes, unicellular cyanobacteria and eucaryotes. While the first two groups are pico-planktonic prokaryotic organisms of specific taxonomic identity, the third group, the eucaryotes are made up by a variety of mostly small celled, flagellated, nano-planktonic species, belonging to diverse taxonomic groupings (Fig. 4.6). Larger eucaryotes, >20 μm in diameter, are also present but usually at low concentrations such that their contribution to total chlorophyll is negligible. 4.5.2.1 The prochlorophytes The prochlorophytes are minute cells, about 0.5 μm in diameter, belonging to a taxonomic group that was first described only in the mid 1980s (Chisholm et al., 1988). Following their initial description, and with the development of epifluorescence and flow cytometric methods that allowed for their quantification, it became apparent that those organisms are very abundant in the world’s oceans, and are the numerically dominant group of phytoplankton in many warm oligotrophic waters at subtropical latitudes. Prochlorophytes contain a unique pigment, divinyl chlorophyll (Goericke & Repeta, 1992), that enables their quantification using HPLC techniques. Prochlorophytes dominate the plankton of the Gulf of Eilat in nutrient-depleted summer stratified conditions (Lindell & Post, 1995). In the Levantine Basin, they are typically vertically stratified with low concentrations in the surface water and peak abundance at the DCM in summer (CYCLOPS, unpublished data), in autumn (Fig. 4.6; Li etal., 1993) and in winter (Zohary etal., 1998). While numerically dominant, their small cell size make their contribution to total chlorophyll only 10–15%. This contrasts to the results obtained at Station ALOHA in the North Pacific subtropical gyre and close to Bermuda in the North Atlantic (BATS) which can be considered as typical oligotrophic regions within the major oceans (Karl & Lukas, 1996; Michaels & Knap, 1996). At these locations, prochlorophytes as determined by divinyl-chlorophyll a accounted for 40–70% (average 57%) of the total chlorophyll a, or about five times the average value recorded for the Eastern Mediterranean (11%) by Vidussi et al. (2001).

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0

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Log10 (Cells ml–1)

Fig. 4.6 Vertical distribution of (䉱) prochlorophytes, (䊊) cyanobacteria, and (ⵧ) eukaryotic ultraphytoplankton at stations across the Eastern Mediterranean (figure from Li et al., 1993).

4.5.2.2 The unicellular cyanobacteria Synechococcus-like unicellular cyanobacteria are the second key component of the pelagic phytoplankton, found in practically all marine environments. Although bigger than the prochlorophytes, the unicellular cyanobacteria are also pico-planktonic measuring less than 2 μm in diameter. In the pelagic waters of the Levantine Basin, they are usually evenly distributed with depth between the surface and the DCM below which their concentrations decline with increasing depth (Li et al., 1993).

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4.5.2.3 The eukaryotes In the earlier taxonomic studies of marine phytoplankton, only eukaryotes were determined, as the prokaryotes were too small to be measured using the available methodologies. Kimor et al. (1987) reported that the most common eukaryotes were coccolithophores, followed (numerically) by diatoms and dinoflagellates. On some occasions silicoflagellates were reported but they were rare. Large species >65 μm in diameter occasionally occurred at low concentrations. An example was the prasinophyte Halosphaera virtidis that was found at the DCM but not in the surface water. Modern HPLC techniques allow characterisation of the eukaryotes by their pigment signatures. HPLC has demonstrated, that prymesiophytes dominated the eukaryotic group in the Cyprus Eddy in March 1992 whereas diatoms were rare. In 2001, dinoflagellates and coccolithophores dominated the eukaryotes (Psarra, personal communication). 4.5.2.4 Heterotrophic bacteria Heterotrophic bacteria are an important component of oligotrophic pelagic environments in general and of the Eastern Mediterranean in particular. In oligotrophic oceans, bacteria often consume up to half the primary production via dissolved organic matter and in turn are consumed by protistan grazers (Azam et al., 1983). Being small in size, usually 0.1–0.5 μm in diameter, which is at the lower end or even smaller than the smallest pelagic phytoplankton cells (the prochlorophytes), bacteria have a large surface area to volume ratio which gives them a competitive advantage in competition for dissolved nutrients. With the present lack of synoptic tools for surveying bacterial abundance, data on the abundance and production of heterotrophic bacteria in the Eastern Mediterranean are more sporadic than for phytoplankton chlorophyll and depend on discrete depth sampling and analysis. The data available are for the Cyprus Eddy in summer (Zohary & Robarts, 1992) and in winter (Zohary et al., 1998), and in the Cretan Sea in March and September (van Wambeke et al., 2000), and a general survey of the Levantine Basin in fall (Robarts et al., 1996). Some additional data exists for the Western Mediterranean which will not be reviewed here. Based on the above publications, the following general features emerge; like chlorophyll, bacterial biomass in the euphotic zone of the Eastern Mediterranean is extremely low at the lower end of abundances reported for oligotrophic oceans. Cho and Azam (1988) suggested a lower threshold for bacterial abundance in the ocean’s euphotic zone of about 3 × 108 bacterial cells L−1. Abundances in the Levantine Basin were consistently below that value, usually ranging between 2 and 3 × 108 bacterial cells L−1 in the upper 100 m. Higher values of 3 to 9 × 108 bacterial cells L−1 were typical of the Cretan Sea. These numbers can be translated to carbon biomass using a conversion factor of 15.6 fg C per cell.

THE EASTERN MEDITERRANEAN

105

Bacterial cells were more or less uniformly distributed with depth between the surface and the DCM. Below the DCM the abundance declined with depth. The very few data points available for deep water of the Eastern Mediterranean (>500 m depth) show that bacterial numbers are at least an order of magnitude lower than at the surface water, although there are occasional “hot spots” of higher bacterial abundance and activity at those great depths. Like chlorophyll, bacterial biomass was also shown to be uniformly distributed horizontally throughout large areas of the Levantine Basin with the exception of mesoscale features where the pattern was usually altered. Probably, the most striking feature of the heterotrophic bacterial population is that its carbon biomass in the euphotic zone is of the same order of magnitude as that of phytoplankton. This is a typical feature of oligotrophic seas stressed by Li et al. (1992). As such, and having a smaller cell size and thus being more numerous in cell abundance than pico-phytoplankton, bacteria are tough competitors for the rare nutrients in these waters. Zohary and Robarts (1998) have demonstrated that bacterial production throughout the Eastern Mediterranean was P limited at least during winter. 4.6

Primary production

Until the mid 1990s measured rates of primary production (14C method) from the pelagic waters of the Eastern Mediterranean were scarce. Mostly, the early measurements suffered from various drawbacks such as being sporadic, using compromised techniques (e.g. on deck or in-lab 14C incubations instead of in situ incubations), taken not far enough offshore, or within enclosed regions rather than at the open sea, and collected at best over a single annual cycle (Becacos-Kontos, 1968, 1977; Becacos-Kontos & Ignatiades, 1970; Oren, 1970; Berman et al., 1984; Dowidar, 1984; Azov, 1986). For example, Azov (1986) claimed that his station 10 km offshore at ca. 250 m depth water was representative of the pelagic waters. We now know from satellite images, such as Fig. 4.7, that on many occasions such coastal waters are not representative of the offshore situation. Primary production values reported by such studies were always at the lower range of values reported for oligotrophic oceans, and the authors used the results to reconfirm the extreme oligotrophy of the Eastern Mediterranean. Recently, it has been realised that these ultra-oligotrophic waters require ultra-clean techniques. This has led to the development of new protocols for the 14 C method, and new data of higher quality and better reproducibility are now becoming available. Most of these new data in the Eastern Mediterranean have been generated by Greek workers in the Aegean and Cretan Seas. The first structured programme to measure primary production at monthly intervals at fixed stations was initiated in 1994 (Psarra et al., 2000) and is ongoing. The

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Jan

Apr

Feb

May

Mar

Jun

July

Oct

Aug

Nov

Sept

Dec

Fig. 4.7 SeaWiFS 9-km monthly composite chlorophyll images for 2000; chlorophyll values in mg m−3 shown on the colour scale. SeaWiFS images courtesy of NASA SeaWiFS Project and Orbimage Inc. (see Color Plate 1).

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Table 4.1 Primary productivity measurements from the Eastern Mediterranean compared with values from other known oligotrophic systems Location Cretan Sea (shelf area) Cretan Sea (slope area) Northwest Sargasso Sea (Bermuda JGOFS site) Northeast Pacific (vertex) Northeast Pacific (vertex)

Primary productivity (g C m−2 y−1)

Reference

80 ± 22 59 ± 22 125

Psarra et al. (2000) Psarra et al. (2000) Lohrenz et al. (1992)

130 120

Martin et al. (1987) Knauer et al. (1990)

two stations being monitored are over the continental slope and the shelf of the Cretan Sea, north of Heraklion. Rates of primary production measured during this programme are generally somewhat higher than those reported in the earlier studies. Averaged over the year they were 80 ± 22 g C m−2 y−1 at the shelf and 59 ± 22 g C m−2 y−1 at the slope station with large inter-annual variability. Lower values (~20 g C m−2 y−1) were reported by Ignatiades (1998), also from the Cretan Sea but from a more pelagic location. The values of Psarra et al. (2000) match well an average winter value of 98 g C m−2 y−1 computed for the Eastern Mediterranean by Vidussi et al. (2001) based on measured chlorophyll and using the spectral photosynthesis model of Morel. These more recent values confirm the unusually low primary productivity of the Eastern Mediterranean compared to other oligotrophic regions of the world’s oceans (Table 4.1). An ongoing routine programme for the Levantine Basin, similar to the one at the Cretan Sea, does not exist and is particularly needed. 4.6.1

Gradient in biomass and productivity from coastal waters to the open sea

This general pattern of species composition in the pelagic Eastern Mediterranean changes in coastal regions (and other situations) where the nutrient inputs to the surface waters are higher. As evident from the satellite images of chlorophyll distribution (Fig. 4.7), several studies document a steep gradient in chlorophyll concentrations and primary production between the coastal waters and the open sea (Berman et al., 1986; Psarra et al., 2000). The coastal phytoplankton populations differ from the pelagic ones both in their size structure and species composition. At such coastal sites, microplanktonic species (>20 μm in diameter) become a significant component of the plankton. Where the added nutrients lead to a bloom, the dominant species is often a large diatom or dinoflagellate. For example, in the coastal region of the Cretan Sea immediately north of Heraklion during 1995, the silica content decreased from 2 μmol kg−1 in winter to less than half that value in summer (Tselepides et al., 2000). This was due to

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water upwelling as a result of the cyclonic eddy that periodically affects this area. It was found that diatoms were a significant part of the photosynthetic community at that time. Similarly it has been found that diatoms can be a significant part of the phytoplankton community in Israeli coastal waters during the winter bloom. By far, the most important example of this effect was the major bloom of diatoms which occurred every summer in Egyptian coastal waters (Halim, 1991) as a result of the nutrient rich flood waters of the Nile being jetted into the Levantine Basin. However this ceased in 1965 when the Aswan Dam was completed.

4.7

Effects of mesoscale features on nutrient and chlorophyll distribution and phytoplankton productivity

As has been noted above, the Eastern Mediterranean is characterised by many eddies and jets (POEM, 1992). Indeed there are almost no areas of the basin which are not part of some mesoscale feature or other (Fig. 4.3). Yet the nutrient distribution (Kress & Herut, 2001) and many of the plankton features such as bacterial abundance and activity and chlorophyll content (Yacobi et al., 1995) seem to be nearly constant across large parts of the basin except for those locations where they intersect major and persistent mesoscale features (Fig. 4.5). Under those circumstances major changes in nutrient distribution and productivity can be seen. The Rhodes Gyre and the Cyprus Eddy (aka Shikmona Gyre) are permanent features which always have an effect on the local biogeochemistry and have been studied in some detail.

4.7.1

Biogeochemical processes in mesoscale features

4.7.1.1 Rhodes cold-core (cyclonic) eddy While there are a number of cold-core (cyclonic) eddies within the Eastern Mediterranean, by far the most prominent is the permanent feature, southeast of the Island of Rhodes, called the Rhodes Gyre (Figs 4.3 and 4.7). A number of studies have been carried out looking at the physical oceanography of this feature, in part because it is thought to be important in the formation of LIW (i.e. Ozsoy et al., 1989; Sur et al., 1993). There are considerably fewer studies available which consider the biogeochemical processes within this feature (Salihoglu et al., 1990; Edigar & Yilmaz, 1996; Yilmaz & Tugrul, 1998). The Rhodes Gyre is a cold-core eddy surrounded by a series of warm-core features that tend to change intensity, size and even location over time. The basic characteristic of the Rhodes Gyre, as with all other cold-core eddies or rings (Fox & Kester, 1986), is that there is an upwelling of nutrients in the centre which results in increased phytoplankton biomass and primary productivity.

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Table 4.2 Table showing seasonal variations in the phytoplankton biomass due to mesoscale features in the Levantine Basin

Type of feature Winter Rhodes Gyre (cold-core) Finike trough (cold-core) Cilician (warm-core) Cyprus Eddy (warm-core) “Normal” Summer Rhodes Gyre (cold-core) Cilician (warm-core) Cyprus Eddy (warm-core) “Normal”

Depth of permanent nutricline

Depth of deep chlorophyll maximum

Integrated chlorophyll a (mg m−2)

Reference

–

–

44*

Ediger and Yilmaz (1996)

–

–

238

Ediger and Yilmaz (1996)

400–600

–

28

Ediger and Yilmaz (1996)

450

–

75

Krom et al. (1992)

200

–

42

Krom et al. (1992)

50–100

50–100

39

Ediger and Yilmaz (1996)

500–600

75–115

30

Ediger and Yilmaz (1996)

450

120

12–39

Krom et al. (1992)

200

100

12–20

Krom (1995)

The winter values are taken from a year cold enough to induce deep winter mixing. * Probably unusually low as a result of active deep winter mixing.

In winter the pycnocline moves up towards the surface. In a mild winter the pycnocline reaches to within 15–20 m of the surface while in a cold winter such as 1991/92 it disappears entirely and LDW reaches the surface. There is then essentially constant temperature and salinity from the surface to at least 1000 m depth (Yilmaz, personal communication). During such periods of deep winter mixing, the surface waters have nutrient contents which are, for the Eastern Mediterranean, extremely high with phosphate values of 0.16 μmol kg−1, nitrate of 4.7 μmol kg−1 and silicate of 6–7 μmol kg−1 and no nutricline present (Table 4.2). This is an important source of nutrients and productivity to the whole basin. The centre of the Rhodes Gyre is the site of the greatest phytoplankton productivity anywhere in the open Eastern Mediterranean away from coastal point sources of nutrients such as the river Nile. For example, the magnitude of the DCM in July is approximately twice that of the rest of the northern basin (Salihoglu et al., 1990) which itself is higher than that found in the southern Levantine Basin. It is also shallower at 60 m than the DCM found elsewhere in the Levantine Basin. This high chlorophyll a value is also shown clearly by remote sensing (Fig. 4.8), where the Rhodes Gyre has consistently much the highest chlorophyll of any offshore area in the basin.

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Eddy centre: Feb. 1989

Eddy centre: Feb. 1989 0.10

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Eddy centre: Nov. 89 0.00

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Eddy centre: Nov. 89 0.40

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Eddy centre: May 89 0.00 0

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Phosphate

Chlorophyll

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Fig. 4.8 Seasonal nutrient distribution of the Cyprus Eddy (after Krom et al., 1992).

However, there are interannual variations caused by the depth and rate of winter mixing. During the exceptionally cold winter of 1991/92 the 3000 m water column at the core of the Rhodes Gyre was vertically mixed from top to bottom, with homogeneous temperature and salinity profiles. At this time, there was relatively low chlorophyll biomass (~44 mg/m2) and phytoplankton activity (Table 4.2; Edigar & Yilmaz, 1996). This cannot be due to nutrient or light limitation since nutrients are present and there are phytoplankton blooms

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occurring at this time in other regions in the northern Levantine Basin. While there have been no direct measurements of vertical mixing rates at this time, it is likely that they are considerable, and that this may be the reason for shortterm limitation of phytoplankton activity. Evidence to support this comes also from the dissolved oxygen profile which is constant to at least 1000 m (the maximum depth sampled) at 210 μmol kg−1, which is the surface (oxygen saturation value) and much greater than the usual 175–180 μmol kg−1 found in LDW (Yilmaz & Turgul, 1998). The nutrient content of the core of the Rhodes Gyre is, however, somewhat lower than that found in the adjacent LDW (by 0.03 μmol kg−1 for phosphate, 0.8 μmol kg−1 for nitrate and 1.5 μmol kg−1 for silicate) presumably due to the mixing down of nutrient depleted surface waters during deep winter mixing. In the periphery of the Rhodes Gyre, where the deep winter mixing was less intense, the phytoplankton biomass (in terms of chlorophyll a and particulate organic carbon) and the primary production rate measurements show higher values which are comparable to western concentrations (Ediger & Yilmaz, 1996). Very high values of chlorophyll a (238 mg m−2) were observed in a less intense cyclonic feature in the Finike Basin at this time. By contrast, in the less cold winter of 1994–95 the core of the gyre only mixed down to 800 m (Malanotte-Rizzoli et al., 1996). In that winter, Vidussi et al. (2001) calculated that the depth integrated chlorophyll content of the 0–230 m layer in the Rhodes Gyre was 61.3 g chlorophyll m−2 which was twice the value of the productivity later averaged for the rest of the basin. The seasonal thermocline is formed in March during mild winters and somewhat later in cold winters. It is generally found at 25–50 m depth. Once the seasonal thermocline has developed, deep winter mixing stops. A nutricline develops at ~75 m which is significantly shallower than that found outside the cold-core feature (Salihoglu et al., 1990). Above the nutricline and DCM, phosphate and nitrate are then below detection limits (Yilmaz & Tugrul, 1998). These measurements were made in July and thus it is not known whether phosphate is entirely consumed immediately the phytoplankton bloom occurs and before nitrate is depleted, as is observed elsewhere in the basin. Unusually in the centre of the Rhodes Gyre, there is a considerable decrease in silicate concentration from 6–7 μmol kg−1 in winter during the deep mixing to 2 μmol kg−1 in the surface waters in summer. The consequence of the higher inputs of nutrients is not only that the depth-integrated chlorophyll content within the gyre is much greater than that measured outside the gyre and elsewhere in the Levantine Basin (Yilmaz & Tugrul, 1998; Zohary & Robarts, 1998; Vidussi et al., 2001). The species composition and size structure of the phytoplankton is also different. The neuston collected in the Rhodes Gyre is much higher than that found anywhere else in the northern Levantine Basin except for very close to the coastal regions. This neuston consists of large phytoplankton (diatoms and dinoflagellates) and zooplankton (copepods, pteropods, isopods, gastropod

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larvae, fish larvae and small shrimp). Thus this nutrient replete system grows large phytoplankton including diatoms which consume part of the available silicate in contrast to the situation in most of the rest of the basin which is dominated by pico- and nano-plankton which have no silicate requirement. The ratio of bacterial to algal biomass is low in comparison with outside the gyre location, which means that the relative importance of bacteria is smaller. 4.7.1.2 Cyprus warm-core (anti-cyclonic) eddy The most detailed studies of warm-core eddies have been carried out by Krom and co-workers on the Cyprus Eddy (aka Shikmona Gyre; Krom et al., 1992, 1993; Zohary et al., 1998). It has also been the site of the recent CYCLOPS phosphate addition experiment which was designed to increase understanding of microbiological processes and nutrient cycling in the Eastern Mediterranean (Carbo et al., 2002). The Cyprus Eddy is a quasi-stationary warm-core feature situated south of Cyprus. It has been found both to the east and west of the Eratosthenes Seamount. It seems to remain stationary for several years and then move out probably to the northeast to be replaced with another similar feature from the southwest (Brenner et al., 1990). In winter there is deep mixing in the core of the eddy. The depth of this mixing depends on the severity of the winter and the history of the particular Cyprus Eddy. The eddy that was sampled in 1989 was mixed to a depth of almost 500 m (Krom et al., 1993), while the Cyprus Eddy sampled in 2001 as part of the CYCLOPS experiment was only isothermal to 300m (Carbo etal., 2002), and that in 2002 to 280 m (Krom et al., unpublished data). During the winter of 1992, when there was deep winter mixing to 500 m, it was calculated that the rate of vertical mixing in the core was at least 30 m h−1 based on phytoplankton light-shade adaptation features and cellular chlorophyll fluorescence per cell. During this deep winter mixing, nutrients which have accumulated above the permanent deep nutricline are mixed up into the surface waters. An immediate phytoplankton bloom occurs which results in apparently viable chlorophyll a throughout the mixed layer (Fig. 4.8; Krom et al., 1992). During this period all of the phosphate is consumed while a 0.5–1 μmol kg−1 of nitrate remains in the water column. The silicate content is 1–2 μmol kg−1 and does not change throughout the year (Krom et al., 1992). In the cold winter of 1992, when winter mixing was ~500 m, the microbial populations were evenly distributed between 0 and 500 m at the core of the eddy, as indicated by chlorophyll concentrations, HPLC-determined pigment composition, flow-cytometric analysis of the ultraplankton, direct counts of bacteria and labelled thymidine measurement of bacterial activity. The depth integrated chlorophyll content was 59 mg m−2. This value was more than double the typical late autumn values suggesting a major bloom was occurring (Zohary et al., 1998). A seasonal thermocline which begins to develop in March stops the deep winter mixing. As a result a DCM develops near the base of the euphotic zone

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at a depth of 110–140 m. Above this DCM, phosphate remains at below the analytical detection limit. Nitrate values gradually decrease until, by May, there is no nitrate measurable in surface waters either. However, recent data has shown that there are significant amounts of DOP, and DON in these surface waters in May which represent a potential reservoir for phytoplankton nutrient uptake (Carbo et al., 2002). Below the DCM, any phytoplankton trapped when deep mixing ceased, appeared to die, decompose and their nutrients were released to form a seasonal nutricline. The DCM then gradually decreases in magnitude during the year as particulate matter is exported from the euphotic zone into deeper layers. Most of the exported nutrients appear to accumulate in the layers between the DCM and the permanent nutricline. This process continues until the end of the year, probably December to February, when deep winter mixing occurs again and the cycle repeats itself. 4.7.1.3 Effects of other mesoscale features It has thus been noted that there are long-term permanent eddies which have a major effect on the biogeochemical processes occurring and other features which seem to have little effect on the nutrient and chlorophyll distribution. Between these two extremes there are also cyclonic and anticyclonic gyres, such as those adjacent to the island of Crete, which wax and wane periodically and their impact on the productivity of the system changes as well. Tselepides et al. (2000) showed that periodically the cold-core eddy adjacent to Crete intensifies greatly and draws in intermediate water which is then upwelled into the surface waters. In 1995, this resulted in a major phytoplankton bloom. Chlorophyll values as high as 1.4 μg l−1 were determined during this bloom which contrast with the normal background values of ~0.05 μg l−1. The blooms adjacent to Crete featured diatoms as a major component of the phytoplankton community. The phytoplankton biomass dies away as soon as the cold-core (upwelling) eddy loses its intensity. The effect of such periodic events can also be seen on the benthic community. It has been observed in the Aegean Sea that pulses of organic rich particulates reach the sediment. These organic matter pulses, which are caused by transient upwelling probably caused by cold-core eddies, stimulate bacterial activity in sediment. A succession of benthic organisms exploit this pulse of food which gradually dies back to the sparse fauna characteristic of the region after a period of several years. 4.8

Seasonal changes in phytoplankton biomass as detected by remote sensing

Figure 4.7 shows the SeaWiFS remote sensing images for the Eastern Mediterranean during 2001. SeaWiFS measures the light scattered out of the

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surface layer of the water in a number of narrow wavebands (see http://seawifs. gsfc.nasa.gov/SEAWIFS.html) chosen at regions of high medium and low chlorophyll absorption. The chlorophyll concentration can be retrieved by comparing the different wavebands (usually using algorithms based on band-ratios: O’Reilly et al., 1998). The depth sensed by SeaWiFS depends upon the concentration of chlorophyll and the waveband; for waters with chlorophyll ~0.1 mg m−3, such as those of the Eastern Mediterranean the depth sensed is ~30 m at 490 nm (Morel, 1988). As has been shown above this represents a good estimate for the phytoplankton biomass during the periods of deep winter mixing but is less representative during the spring and summer periods when there is a DCM at 100–140 m across the Eastern Mediterranean. The SeaWiFS composite images show clearly the seasonal productivity succession in the Eastern Mediterranean. The annual phytoplankton bloom starts in October. Initially, productivity develops and expands from the coastal regions, particularly the southern coast of Italy into the Ionian Sea, the northern Aegean into the southern Aegean and the shore of Israel–Lebanon–Syria into the eastern Levantine Basin. Mesoscale features are important in the developing bloom with cold-core eddies being the first to show clear signs of increased phytoplankton biomass. Even in the coastal regions, it seems that eddies and other mesoscale features caused by interactions between the currents and the coast, are important as sites where bloom conditions are initiated. This can be seen particularly clear in the eastern coast of the Levantine Basin. However, it should be noted that coastal waters entrained offshore by mesoscale activity may contain suspended particulate material (SPM) and coloured dissolved organic matter (CDOM) that produce erroneous chlorophyll retrievals. By November there are only rather small areas of the southern and particularly southeastern Levantine Basin that still retain the summer low biomass condition. The widespread and intense phytoplankton bloom lasts in the Eastern Mediterranean from December until mid-March. By March the seasonal thermocline would have reappeared and the characteristic summer conditions begin to reestablish themselves initially in the southern Levantine Basin. With the initiation of the bloom, mesoscale features control the details of the re-establishment of summer conditions. Eddies can still be clearly seen as areas of higher biomass in the southern Levantine, particularly in March. Summer conditions with extremely low productivity in the surface layers, and a DCM is fully established by May and reach their peak between June and August, when all of the southern Levantine has extremely low chlorophyll levels in the surface layers. It is important to remember that the satellite only sees the phytoplankton in the uppermost layers of the system. Once the seasonal thermocline is developed, the system is characterised by a deep chlorophyll maximum. Thus the drastic changes in chlorophyll shown by these images are in reality less marked.

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Superimposed on this dominant seasonal pattern, geographical differences can be seen. The northern Levantine Basin, the Aegean and the northern Ionian Seas have higher chlorophyll levels throughout the year than do the southern Levantine Basin and the southern Ionian Sea. While some of this difference may be due to shallower waters, that is only a small part of the picture since there is water in excess of 3000 m in both the Aegean and northern Ionian Seas. An important feature observed previously are the eddy features on the coast of Israel, Lebanon and Syria (Karabashev et al., 2002). These turbulent features which are presumably related to the general cyclonic circulation of the Mediterranean interacting with the coast may be important as sites where higher productivity is initiated and nutrients exported into the southeastern Levantine Basin. The Rhodes cold-core eddy can be seen at all times of the year as a prominent and permanent feature of increased productivity. While there are other cold-core eddy features within the basin which can be seen on this set of images none of them show up as such a clear and important feature. The most prominent and permanent warm-core eddy is the Mersah Matruh Gyre situated off the coast of Egypt–Libya. It can be seen as a strong low productivity feature in July and August. Interestingly this feature does not show up as a low productivity feature in winter. Weaker features such as the Cretan cold-core eddy and the Cyprus warm-core eddy can be seen as features which show up occasionally. While there is a net export of nutrients at the Straits of Sicily, elevated chlorophyll can clearly be seen from the satellite images, particularly in spring, possibly associated with nutrients brought into the eastern basin from the Western Mediterranean. The major terrestrial input into the basin is the river Po which discharges into the northern Adriatic. The plume from the river flows to the south along the east coast of Italy. There seems to be a seasonality to the discharge with the highest chlorophyll presumably caused by nutrient and/or particulate input occurring in October and November. There is a clear effect observed for the input from the river Nile. Prior to 1965, there was a major input of nutrients into the Levantine Basin during the time of the annual Nile flood which reached the Egyptian coast in August–September. This plume used to flow around the coast of the Levant appearing as a layer of brown high productivity water off the coast of Israel and Lebanon (Hornung, personal communication). The major phytoplankton species in this bloom used to be diatoms. The silica required for this bloom was provided by the silica-rich Nile river water. This is no longer occurring. What can be seen is a relatively constant phytoplankton bloom caused by a constant input of nutrients throughout the year; though as noted above, the chlorophyll retrievals are likely to be affected by the presence of SPM and CDOM. Furthermore the input appears to be from the entire front of the Nile delta with no particular hot spots corresponding to the Damietta and Rosetta mouths of the river. While it has been argued by Nixon

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(2003) that the total amount of nutrients supplied to the basin has been restored to the levels that occurred prior to 1965, he notes that there is now no seasonality to this discharge. It is also likely that the species composition of this annual bloom will have changed because the nutrients which reach the Mediterranean are now higher in nitrate and phosphate and relatively lower in silicate. At this time no detailed studies have been carried out on this major anthropogenic change to the biogeochemistry of the basin. The Black Sea shows up as an area of very high chlorophyll throughout the year. This is expected since it is dominated by river inputs and estuarine circulation. The sea of Marmara also has high chlorophyll throughout the entire year. While there is a higher chlorophyll plume from the Black Sea into the northern Aegean, the amount seems to be relatively small and certainly smaller than the plumes observed from the Po and the Nile. Qualitative observations on the biomass produced by these extra-basin inputs are entirely compatible with recent estimates for the amount of nutrients fluxed into the basin from the sources (Krom et al., in preparation). In addition to these natural increases in biomass there is a major increase in colour off the Tunisian coast. This corresponds to an area of shallow water where there is both an increase in chlorophyll and particulates in the water column.

4.9

Nutrient limitation in the Eastern Mediterranean

In most parts of the world’s oceans, the molar ratio of nitrate:phosphate in the deep water is 16:1, which is also the ratio in which these elements are taken up in phytoplankton growth (Redfield et al., 1963). In the Eastern Mediterranean the N:P ratio in LDW is closer to 29:1 (Krom et al., 1991). This means that when the waters are mixed upwards in winter they have an excess of nitrogen over phosphorus compared with the Redfield ratio. The winter phytoplankton bloom ceases as soon as the system runs out of phosphate. Characteristically this happens when there is 1–2 μmol kg−1 of nitrate remaining in the surface waters. This is the usual definition of nutrient limitation and is one of the strong pieces of evidence used by Krom et al. (1991) when they suggested that the Eastern Mediterranean was P limited. Subsequently it has been found that the δ15N-NO3 content of this residual nitrate is unusually heavy (16–40% compared with 3–12% at depth; Struck et al., 2001). During photosynthesis, light nitrate is used up preferentially to heavy nitrate. If the consumption of nitrate stops because the system is limited by some other nutrient or factor, then this heavy nitrate remains within the system. However, in summer the surface waters of the Eastern Mediterranean above the DCM are entirely nutrient depleted with both nitrate and phosphate absent from the surface waters (Krom et al., 1992; Tselepides et al., 2000; Kress & Herut, 2001). This would imply that the system had become co-limited by

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nitrate and phosphate. At present there is no explanation for this unusual feature of the Eastern Mediterranean system. Additional evidence for P limitation of the surface waters of the Eastern Mediterranean has been obtained from observations on phytoplankton and bacterial activity. Pojed and Kveder (1977) working in the northern Adriatic believed P was the limiting nutrient for the growth of phytoplankton. Vukudin and Stojanski (1976) were of the same opinion from their work in the central and southern Adriatic. Becacos-Kontos (1977) suggested that the factors limiting primary productivity in the waters of the Gulf of Saronique (Aegean Sea) were sometimes solely P and sometimes both P and N. Bonin et al. (1989) showed that C and P uptake by natural microplankton populations collected from two stations adjacent to the coast of Israel was enhanced by adding phosphate, or phosphate and nitrate, but not by adding nitrate alone. Similar results have been reported farther offshore in the Levantine Basin by Berland et al. (1987). Later, Berland et al. (1990) reported that water from the middle of the Levantine Basin had high particulate organic carbon (POC)/particulate organic phosphorus (POP) and particulate organic nitrogen (PON)/POP ratios, which they interpreted as indicating a severe phosphorus limitation of algal biomass and specific division rates. However, the evidence is not unequivocal since they found that a nitrogen supplement was also needed for any obvious increase in biomass. Recently, Thingstad et al. (2001) found that productivity was enhanced in microcosms to which phosphate alone or phosphate and nitrate were added, but they also found that nitrate alone increased productivity. Zohary and Robarts (1998) showed that in winter, bacteria from various sites in the Eastern Mediterranean grew when P was added, whether N was added or not, but did not grow faster than controls when N alone was added, or even when N + iron + EDTA (as chelator) were added. Experimental evidence is gradually accumulating, demonstrating that at least during part of the year P was the limiting nutrient for algal and bacterial growth (Zweifel et al., 1993; Berdalet et al., 1996; Vaulot et al., 1996). It seems likely that while the Eastern Mediterranean is P limited in winter, the evidence in summer is equivocal. Furthermore, it is possible that bacteria and phytoplankton may not have the same nutrient limitation behaviour in this system. It is expected that new insights into the nutrient limitation of the Eastern Mediterranean will arise when the results from the CYCLOPS programme are to be published in 2003 and 2004. However, it is clear that the Eastern Mediterranean is not iron limited. Those areas of the ocean which are iron limited are characterised by high nitrate and phosphate remaining in the surface water when there is sufficient light to allow phytoplankton growth and yet productivity has stopped. Both the East Central Pacific and the Southern Ocean are thousands of kilometres downwind of the nearest source of atmospheric dust and far away from any land source of iron. Recent results of the IRONEX and SOIREE lagrangian

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experiments have shown that under these circumstances the system becomes iron limited. The Eastern Mediterranean has none of these characteristics. It has no residual nutrients left in the surface water in summer. It has one of the highest fluxes of dust of any ocean in the world (Chester et al., 1977; Guerzoni et al., 1999). Furthermore preliminary measurements of dissolved iron in the Eastern Mediterranean by Mantoura et al. (personal communication) showed values of 2–4 nmol kg−1 which are far above the levels at which iron has been found to be limiting (Martin et al., 1987). Of more relevance is the suggestion that N fixing phytoplankton require a large amount of iron for their specific enzyme requirements, and may be stimulated by iron. In systems where there is a large excess of iron supplied by dust, it has been suggested that there is significant N fixation even when inorganic N and P are absent (Wu et al., 2000). Such N fixation has been suggested as the reason why the N:P ratio in the deep water of the Eastern Mediterranean is >>16:1 (Bethoux et al., 1999). At this time, no occurrence of Trichodesmium, which is the organism generally thought to be important in such planktonic N fixation (Karl et al., 1997), has been observed in the Eastern Mediterranean. While Synechococcus is common in the surface waters of the Eastern Mediterranean, there is limited evidence that this species of cyanobacteria actually fixes gaseous nitrogen. Specific experiments to investigate N fixation in the Eastern Mediterranean have been carried out within the CYCLOPS programme. 4.10

Magnitude of man-induced changes in nutrient inputs and their possible effects on the Eastern Mediterranean

Nutrient budgets for the Eastern Mediterranean have been carried out by MacGill (1969), Bethoux (1981), Sarmiento et al. (1988), Vollenwieder et al. (1996) and Bethoux et al. (1992, 1998), amongst others (Table 4.3). The most recent of these estimates by Krom and co-workers uses the seasonality of the flow data of Astraldi et al. (1999) to create a more accurate budget. These budgets show the magnitude of the net export of nutrients from the basin which

Table 4.3 Calculated net flux of nitrate and phosphate out of the Eastern Mediterranean through the Straits of Sicily Nitrate

Phosphate

Reference

142 170 230

5.0 7.2 9.8 1.7

Krom et al. (in preparation) Bethoux et al. (1998) Bethoux et al. (1998) Sarmiento et al. (1988)

Fluxes are given in 109 mol y−1.

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Table 4.4 Table of the estimated atmospheric and terrestrial input (ATI) to the Eastern Mediterranean compared to estimates for the anthropogenic input of nutrients to the basin Estimates

Nitrogen

Phosphorus

Reference

Estimated ATI for eastern basin Estimated ATI for eastern basin Anthropogenic input (UNEP scenario 2) Anthropogenic input (UNEP scenario 1)

142 230 63

5.0 9.8 3.3

Krom et al. (in preparation) Bethoux et al. (1998) UNEP (1996)

13

0.8

UNEP (1996)

(All values are given in units of 109 mol y−1).

cause the waters to the east to be dominantly oligotrophic despite being surrounded by land and subject to significant natural and anthropogenic nutrient input. The total net flux of nutrients (N and P) through the straits of Sicily has been estimated as 142 and 5 (109 mol y−1, Table 4.3). At steady state, this net flux is equal to the input of nutrients into the basin from atmospheric and terrestrial inputs such as continental weathering, wet and dry atmospheric input, nitrogen fixation and direct pollution input, once correction has been made for losses to the sediment (Bethoux et al., 1999). The assumption of steady state depends to some degree of the time and space scale that is being considered. Bethoux et al. (1998) argue that there has been a steady increase of 0.5% per year in the nitrate and phosphate content of the deep waters of the Western Mediterranean caused by a 3% annual increase in anthropogenic flux to that basin. Others have argued that no such systematic trend can be observed in the existing data sets (Denis-Karafistan et al., 1998). Nobody at present has asked the same question from the Eastern Mediterranean data. This is partly because the deep basin has a much higher volume and the turnover time of the deep water is longer which means that recent anthropogenic changes would take longer to effect the system. In addition, there are few long-term nutrient data sets which could be used to examine this question. However, there is no doubt that the system has been subject to major anthropogenic stress, and it is reasonable to assume that as a result the nutrient flux has changed in both magnitude and nature. Estimates have been made of the anthropogenic discharges into the basin using two different scenarios (Table 4.4; UNEP, 1996). In the first scenario, the N and P input from the human population in the coastal zone was estimated assuming it reaches the basin as sewage in one form or another. In the second scenario, the N and P discharged into the sea from various human activities including agriculture, food processing and sewage was estimated. These inputs represent up to 40% of the N and 66% of the P influxing into the basin. Furthermore of the remaining input to the basin, which includes atmospheric and

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riverine discharges, it is far from clear what fraction is “natural” or indeed what the term natural means in the context of a basin where civilisation has been actively present for the past 5000 years or more. In recent times, the discharge of almost all the rivers flowing into the basin has been drastically changed and generally reduced, with the most dramatic example being the river Nile, which lost 90% of its natural flow as a result of the completion of the Aswan Dam in 1965. Krom et al. (in preparation) have calculated that the major input of nutrients to the basin comes from atmospheric input. Again it is almost impossible to estimate how much of this input is “natural”. There clearly has been a significant long-term increase in the N content in the atmosphere as a result of the burning of fossil fuels, particularly from cars. The P content has probably also increased recently as a result of agricultural practices.

4.11

Summary and conclusions

The Eastern Mediterranean is an unusual and unique marine ecosystem. It is almost landlocked, yet despite receiving considerable inputs of natural and anthropogenic nutrients from atmospheric and landbased sources, it is ultraoligotrophic. It has very low depth integrated chlorophyll, low annual primary productivity, a summer deep chlorophyll maximum deeper than 100 m and a high content of pico- and nano-plankton. This is due to the low level of nutrients in the water column where phosphate reaches a maximum of <0.3 μmol kg−1 and nitrate of ~6 μmol kg−1 in the deep water. The reasons for this low nutrient concentrations and ultra-oligotrophic water is that nutrients are exported from the system at the Straits of Sicily by the anti-estuarine circulation in the system. The Eastern Mediterranean is also unusual in being the largest body of water in the world in which the primary productivity is phosphorus limited. The annual phytoplankton bloom occurs in winter as soon as deep mixing occurs and ceases when the phosphate is depleted. The nitrate which remains is enriched in heavy 15N-NO3. Nutrient enrichment experiments of bacteria show P limitation throughout the year. There is no evidence of Fe limitation in this system, and thus far no evidence of significant N-fixation by phytoplankton has been found. After the winter bloom, water column stratification occurs in March–April, which results in the formation of the DCM that is characteristic of the system for the remainder of the year. The stratification starts in the offshore area in the Southeast Levantine Basin and spreads from there to the north and west. In summer the DCM is typically greater than 100 m deep with the waters above it depleted in inorganic N and P but containing significant amounts of dissolved organic nitrogen and phosphorus.

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The phytoplankton community in the offshore region is dominated by pico- and nano-plankton, with microplankton and eukaryotes only important in coastal regions and upwelling areas associated with mesoscale features such as the Rhodes Gyre. Heterotrophic bacteria are an important component of this nutrient depleted system. Both depth integrated chlorophyll a and primary productivity levels are extremely low, characteristic of ultra-oligotrophic systems. Mesoscale features can be locally important in controlling biogeochemical processes. The Rhodes cold-core gyre is a locus of nutrient upwelling and an important source of both nutrients and productivity for the whole basin. It is also the major source of LIW. The Cyprus warm-core eddy, which has been studied extensively, has elevated productivity in winter as nutrients are brought to the euphotic zone by deep winter mixing. The depth integrated chlorophyll decreases during the year but never becomes as low as the area around the eddy. Satellite images show the importance of transient mesoscale features as loci for seasonal biomass changes and for nutrient transfer. Although it is known that anthropogenic inputs of nutrients are important and possibly dominant in the overall supply to the system, there are no long-term nutrient data sets available to monitor changes in nutrient status with time, as there are in the Western Mediterranean. It is however known that there have been recent changes in the deep water circulation in the Eastern Mediterranean which are probably caused by climatic and/or anthropogenic changes. The Eastern Mediterranean is extremely sensitive to such changes. Only 5000 years ago, the basin was like the Black Sea with an oxygen depleted deep water. We are still unsure of the exact conditions which caused this major shift from an ultraoligotrophic to a eutrophic system though increases in nutrient supply and paleoclimatic changes are prime candidates.

Acknowledgements The authors would like to acknowledge the helpful and lively discussions with the CYCLOPS partnership including Fauzi Mantoura, Frede Thingsted, Barak Herut, George Zodiatis, Malcolm Woodward, Cliff Law and Tassos Tselepides, which took place in the dry laboratory during the cruise of the RV Aegeo to the Southeast Levantine Basin in 2002 and which helped develop the ideas in this chapter. We would like to thank Vyvyan Codd for ordering the raw satellite data and data processing. It was written as part of the background work for the CYCLOPS programme (EVK3–1999–00037) though unfortunately was completed before the results of the CYCLOPS Lagrangian experiment in the Eastern Mediterranean were available for publication. It was completed while MDK was on sabbatical leave, funded by the Leverhulme Trust (RFG/ 10307).

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Glossary AVHRR NASA NOAA SeaWiFS SST

Advanced Very High Resolution Radiometer National Aeronautics and Space Administration National Oceanic and Atmospheric Administration Sea-viewing Wide Field-of-view Sensor Sea-Surface Temperature

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5

The Arctic seas Michael L. Carroll and JoLynn Carroll

5.1

Summary

The Arctic Ocean, together with its associated marginal seas, is one of the smallest of the world’s ocean areas. However, the geographical and environmental peculiarities of the region result in the Arctic having a disproportionately large influence on the world ocean system. The presence of permanent and seasonal sea ice, as well as the seasonal input of large volumes of freshwater and associated constituents from rivers, have a profound influence on the character of seawater and biogeochemistry of the Arctic marine system. The Arctic Ocean strongly influences the global ocean circulation through thermohaline circulation and deep water formation. Strong allochthonous inputs of nutrients from rivers and major ocean currents flowing northward from mid-latitudes as well as from deep water reservoirs, combined with stratification from spring sea ice-melt, lead to localized areas of high primary productivity in the Arctic marginal seas. The great extent of shallow shelf seas (35% of total Arctic Ocean area) further leads to vigorous transformations of organic matter on the continental shelves with off-shelf export to the deeper waters of the interior Arctic basin. Species diversity tends to be lower and food chain length shorter in the Arctic compared to lower latitude oceans, reflecting the rigorous environmental conditions and specialized adaptations necessary to thrive in the Arctic. However, the seasonal localized pockets of intense primary production and high biomass of lower trophic level herbivores often support an abundance of higher trophic level organisms (birds, seals, whales), many of which migrate seasonally into the Arctic to feed and reproduce. The Arctic ecosystems’ limited biodiversity, short food chains and specialized adaptations to the local environment (e.g. lipid storage) render them particularly susceptible to environmental perturbations. Contamination from either local or far-afield sources, long-term climatic changes and variations in ultraviolet-B radiation (UVB) are the most likely factors to significantly impact arctic ecosystems. It has been predicted that the region will experience amplified effects of global climatic changes. The possible consequences to arctic ecosystems range from negligible to profound. Although primary production may increase in certain areas as a result of higher temperatures and less sea ice, changes in the timing and location of the spring bloom may lead to displacement

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of species and/or communities, with as yet unknown consequences to biological resources that are currently exploited (i.e. fisheries). A long-term warming trend may have acute negative consequences to the specialized obligate sea-ice associated communities. From a global perspective, deep water formation in the Greenland Sea – a process thought to be instrumental in driving global ocean circulation patterns – has a tremendous influence on the weather and climate patterns of the world, especially heavily populated continental Europe. The stability of such a process has come into question under some scenarios of climate change. For such an extreme environmental setting, the Arctic has a surprising abundance of natural resources and provides a variety of ecosystem services to users ranging from local residents to the global population. Cascading ecosystem effects resulting in changes to the distribution and abundance of middle and higher trophic level species can have significant socio-economic consequences to indigenous and non-indigenous residents of the Arctic. 5.2

Main features

The Arctic Ocean and its bordering seas cover an area of 14 × 106 km2. It is comprised of a deep central basin (>5000 m) bisected by the Lomonosov Ridge, separating it into the Eurasian Basin on the Atlantic side and the larger Canadian Basin on the Pacific side. The Arctic Ocean is connected to the Pacific Ocean via the shallow (<100 m) Chukchi Sea and Bering Strait. Connections to the Atlantic Ocean are via the Fram Strait, the Norwegian and Barents Seas, and through the Canadian Archipelago and Baffin Bay. With the exception of Fram Strait, water from both the Pacific and Atlantic sides transits a shallow continental shelf on its way into the interior Arctic (Fig. 5.1). While inflows occur from both the Pacific and Atlantic, outflow occurs only on the Atlantic side. 5.2.1

Water masses

The circulation patterns and distribution of water masses in the Arctic are a fundamental element in the biogeochemistry of the Arctic seas. Most of the Arctic Ocean water originates in the North Atlantic, which provides almost all of the Arctic’s mid and deep water, and is the primary source of oceanic heat to the Arctic Ocean. Water enters from the North Atlantic via two pathways: through the eastern Fram Strait via the West Spitsbergen Current (1–1.5 Sv, 1 Sv = Sverdrup = 1 × 106 m3 s−1) (Rudels et al., 1996; Schauer et al., 1997; Rudels & Friedrich, 2000) and via the Barents Sea Overflow (2 Sv) whereupon it is cooled and freshened (Steele et al., 1995; Schauer et al., 1997) (Fig. 5.2). In winter there is a large heat loss from this inflowing water, and brine is injected into the underlying waters as sea ice forms. These processes add enough

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it stra ing Ber

Legend Land 1000 m isobath 4000 m isobath 5000 m isobath major river inflow

ALASKA

EAST SIBERIAN SEA

CHUKCHI SEA

LAPTEV SEA

Canadian Basin

CANADA

ge

v

so

o on m o

L

Ba

Eurasian Basin

RUSSIA St. Anna Trough KARA SEA

trait

GREENLAND

Fram S

y

in

ff Ba

id R

GREENLAND SEA

Svalbard

BARENTS SEA

NORWAY

NORWEGIAN ICELAND SEA

Fig. 5.1 Circumpolar view of the Arctic Ocean showing bathymetry, major surface current (heavy arrows) and bottom current (thin arrows) patterns, and locations of major river inflows.

density to the water masses to allow them to flow off the continental shelves in plumes, entraining and mixing with other waters as they sink (Rudels et al., 1994; Jones et al., 1995; Jones, 2001). The Pacific water entering through the Bering Strait and the Chukchi Shelf (0.8 Sv) is relatively cold and has low salinity, and contributes mainly to surface layers in the Canadian Basin. River discharge is also a major contributor to Arctic circulation and water mass composition (Table 5.1). The Arctic seas receive approximately 4200 km3 y−1 (AMAP, 1997), about 10% of the global river runoff (Aagaard & Carmack, 1989). This discharge is central in the establishment of the Arctic Ocean’s halocline (Aagaard et al., 1981; Steele et al., 1995; Rudels et al., 1996; Schauer et al., 1997) and in the interhemispheric transport of freshwater (Wijffels et al., 1992). Cooling, salt distillation effects associated with ice formation and melting, and mixing physically modifies the various inflows crossing the Arctic marginal seas before entering the interior basin. After transiting the Arctic, water is entrained in the Transpolar Drift moving toward Fram Strait, or in the

SHELF

BEAUFORT WINTER WATER

CANADA BASIN

THERMOCLINE

LOWER PYCNOCLINE

UPPER PYCNOCLINE

MACKENZIE R.

BEAUFORT SEA

YUKON R.

KOLYMA R.

L R

KARA SEA

YENISEI R.

WSC

ICE MELT

BSO

BARENTS SEA

EURASIAN BASIN

SHELF DRAINAGE

LAPTEV SEA

LENA R.

ATLANTIC LAYER

BERING INFLOW

BSI

INDIGIRKA R.

CHUKCHI EAST SEA SIBERIAN SEA

RUSSIA

OB R.

OUT

FRAM STR.

IN

L R

GREENLAND

EGC

GREENLAND SEA

CANADA BASIN

BA F BA FIN Y

Fig. 5.2 Schematic of the circulation and stratification patterns associated with river inflow and halocline ventilation in the Arctic Ocean. (a) Surface view of the Arctic Ocean showing the location of major inflow and outflow currents and relative magnitude of inputs of major rivers. (b) Hypothetical section following the shelf break around the Arctic basin and various water mass strata within the ocean. LR = Lomonosov Ridge; EGC = East Greenland Current; WSC = West Spitsbergen Current; BSO = Barents Sea Overflow; BSI = Bering Sea Inflow. Modified from Carmack (1990). Reproduced with Permission, Academic Press 1990.

(b)

(a)

CANADA

SEV DVINA R.

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Table 5.1 Annual freshwater discharge of major rivers to the Arctic sea River

Discharge (km3)

Yukon Mackenzie Nelson Northern Dvina Pechora Ob Yenisey Lena Kolyma All other discharge

210 333 75 106 140 404 630 525 132 1645

Source: AMAP, 1997.

anti-cyclonic Beaufort Gyre in the Canadian Basin. The relative strength of these two current paths varies strongly with wind fields (Dickson et al., 2000; Maslowski et al., 2000). Outflow of near-surface Arctic water occurs through the Canadian Archipelago (1 Sv) exiting through the Labrador Sea, while deeper water exits the Arctic through Fram Strait via the East Greenland Current (Fig. 5.1). The vertical structure of water masses in the Arctic Ocean is characterized by a surface polar mixed layer extending to a depth of 50 m (Fig. 5.2). The surface layer is comprised of cold, low salinity water and is isolated by a distinct cold halocline from the underlying warm Atlantic water layer (approximately 3°C; 200–700 m depth range). The genesis of the halocline lies in the West Spitsbergen Current, which is cooled and freshened upon encountering the ice edge (Rudels et al., 1996). The inflowing water progresses eastward along the continental shelf, undergoing freezing and thawing cycles with convection, down to the warm Atlantic layer. Freshwater inflow from the Laptev Sea ultimately provides a cap on this convection (Jones, 2001) establishing the upper and lower haloclines in the Canadian Basin. Below the Atlantic layer are the upper polar deep water (to 1700 m), deep water (to 2500 m) and bottom waters. The Greenland Sea east and south of the Fram Strait is a key oceanographic area where overturning of cold polar surface water in winter months results in subsidence to great depths (Davis, 2000). This deep convection renews deep water, eventually becoming Atlantic bottom water, forming a major initiating mechanism in global thermohaline circulation (Fig. 5.2). 5.2.2

Continental shelves

While the Arctic Ocean represents about 2% of the global ocean volume, its associated continental shelves represent about 25% of the global shelf area.

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These statistics underscore the importance of the continental shelf areas to the structure and function of the Arctic Ocean system. Arctic shelf seas are the primary sites for processing and modifying the characteristics of waters received from the Pacific and Atlantic Oceans and the numerous large rivers that drain the circumpolar continents. All of these inflows are substantially altered on the shelves by mixing and by interactions with the ice cover, atmosphere, seabed and biota. The water mass properties that generate and maintain the halocline in the Arctic Ocean (see Section 5.2.1) are derived from the modification of inflowing Atlantic and Pacific waters while transiting continental shelves. Continental shelves are the primary locations for annual ice formation and melting, and are the locations of the highest primary productivity in the Arctic. The large continental shelves of the Arctic are also important for transporting atmospheric CO2 to deeper regions of the ocean (Anderson, 1995; Anderson et al., 1998). Arctic shelf and slope sediments contain a significant reservoir of methane. Increasing Arctic temperatures may liberate methane to the atmosphere and subsequently contribute to the greenhouse effect. These areas, thus play a central role in regulating many of the physical and biogeochemical balances that maintain the arctic natural system with each of the shelf sea areas exerting a unique influence on the biogeochemical composition of the interior Arctic Ocean. 5.2.3

Sea ice

Sea ice is perhaps the defining feature of the polar seas. At its annual maximum extent, it covers about 13% of the earth’s surface area (Thomas & Dieckmann, 2002). Sea ice extent in the Arctic varies from a winter maximum of approximately 15 × 106 km2 to a minimum in summer of 8 × 106 km2. Thus about 40% of Arctic sea ice is seasonal. Its presence directly and indirectly mediates many of the physical, chemical and biological processes of Arctic systems. Sea ice acts as an effective insulator, diminishing the heat fluxes between atmosphere and ocean, and significantly influences the albedo of the ocean surface. Sea ice also represents a physical habitat for an ice-associated community and is a large reservoir of freshwater. The distribution of the permanent ice pack is due to the Arctic Ocean’s halocline. This strongly stratified layer inhibits the vertical flux of heat into the surface layer from the vast pool of relatively warm water found at mid-depths in the Arctic Ocean. Annual or first-year ice is generally formed in the latitudinal zone between 60 and 75° N. At the end of summer, sub-freezing air temperatures cool the surface water. When the surface sea water reaches −1.8°C, ice begins to form. During the freezing process, salt is ejected as concentrated brine, accumulating in grooves and brine pockets, and is transferred to the surrounding seawater. Sea ice is in constant motion. Fast ice located along coastal boundaries is made relatively stable by anchoring to the shoreline and to the sea bottom

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in shallow areas. The width of the fast-ice zone varies from a few kilometres (Point Barrow, Alaska) to over 100 km (the shallow Siberian shelf) (Gow & Tucker, 1990). Divergent stresses of winds and currents cause cracks in the ice surface which can widen to form leads. These areas of open water range from metres to kilometres wide and are a common feature of Arctic sea ice. The locations of leads are generally unpredictable and these features are largely ephemeral, closing again following a shift in wind or by rapidly refreezing during winter (Davis, 2000), a process which can result in ice ridges several metres high. In contrast to leads, polynyas are regions that remain ice-free in predictable and recurrent locations and at times of the year when climatic and oceanic conditions would normally dictate ice cover. Two different physical mechanisms are responsible for the formation of polynyas (Muench, 1990). Latent-heat polynyas occur when sea ice does form, but is continually removed by recurrent wind or sea currents. This type of polynya is often associated with a coastal feature such as a coastline or channel constriction that restricts a compensatory flow of ice. Sensible-heat polynyas, in contrast, are maintained in areas where oceanic heat (i.e. warmer waters) enters the surface region in sufficient quantity to prevent ice formation locally.

5.3

Biogeochemical cycles and ecological processes

High latitude pelagic ecosystems are characterized by strong seasonal cycles driven by the availability of light and inorganic nutrients, the necessary prerequisites for photosynthesis. The primary chemical precursors to biological production (carbon, oxygen and the essential nutrients) are combined by autotrophs to form organic matter (production). When organic matter is broken down into its component elements (decay) it is either recycled within the water column or permanently buried under the seafloor. The biogeochemical cycles of the precursor elements (C, N, P, O2) as well as silica (Si), sulphur (S), barium (Ba), and iron (Fe) have essential roles in the bioproduction of the oceans. Although oversimplifying the complexities of organic matter formation, the ‘Redfield Ratio’ (C/N/P = 106/16/1) (Redfield et al., 1963) is supported by measurements in most areas of the world’s oceans including the Arctic Ocean (Anderson & Dyrssen, 1981). Deviations from the Redfield Ratio are indicated in some areas, for example in the shallow Barents Sea where higher C/N ratios (~8.75) are required to balance the chemical composition of these waters (Kaltin et al., 2002). Superimpose the attributes of the Arctic environment upon the cycles of these biological precursors, and what arises is a system of biological production, transformation and losses unique in the global ocean system. Continental shelves are a key feature that influence greatly the production and cycling of organic matter in the Arctic Ocean. Water masses exiting different Arctic shelves

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bear the distinct imprint of individual shelf regions. Pacific waters flowing through the Bering Strait accumulate products of organic matter breakdown from the water column and/or sediments during transport across the shallow Bering Sea shelf and lead to a pronounced nutrient maximum in the upper water of the Canadian Basin (Codispoti & Owens, 1975; Cooper et al., 1997). The Arctic Ocean as a whole is noted for having the highest concentrations of dissolved organic carbon (DOC) of all major ocean basins (Anderson et al., 1994; Bussman & Kattner, 2000). In highly productive Arctic shelf areas, decoupling between water column production and consumption is thought to control significant offshore transport of particulate organic carbon (POC) as well as DOC (Walsh et al., 1989, 1997; Walsh, 1995). Outflows driven by physical processes effectively transfer organic material from the shelves into the subsurface layers of the Arctic Ocean (Björk, 1989; Guay et al., 1999; Fransson et al., 2001). Halocline formation and shelf-basin exchange may be the major physical mechanism supplying organic material to the basin (Jones et al., 1990), since in situ production is so limited. At high latitudes, the microbial food web significantly influences the vertical export of primary production (Reigstad, 2000; Olli et al., 2001). Bacterial growth efficiency is related to temperature, leading to the export of a larger proportion of primary production from the upper ocean to depth in polar regions as compared to lower latitudes (Rivken & Legendre, 2001). However DOM cycling on Arctic shelves and in the basin is different. Within the permanently ice-covered central Arctic, relatively high ratios of bacterial activity to primary production are observed indicating net heterotrophy and a sink for DOM (Rich et al., 1997). In contrast, ratios of bacterial production to primary production observed in the Chukchi Sea are relatively low (Cota et al., 1996) supporting the view that the marginal Arctic seas are potential exporters of DOM to the interior Arctic basin. Terrestrial and marine biomarkers have aided in understanding how Arctic shelf systems process, metabolize and sequester carbon (Opsahl et al., 1999). Biomarkers have been used to trace the dispersal of DOC and POC on the Canadian Beaufort shelf, leading to a better understanding of the importance of terrestrial sources to the carbon budget of major riverine-influenced systems (Yunker et al., 1995; Macdonald et al., 1998). The biogeochemistry of the Arctic Ocean is markedly influenced by the growth, distribution, and decay of organisms, particularly those in lower trophic levels comprising the largest component of biomass. Although much of the Arctic Ocean is relatively unproductive, limited by a lack of solar energy during much of the year, nutrient availability and extensive sea ice, high biological production occurs in certain locations of the Arctic Ocean during the summers, mainly in the shelf seas (Walsh et al., 1989; Sakshaug et al., 1994; Grebmeier et al., 1995). These highly productive areas in turn supply the central Arctic basin with DOC and POC, nutrients, sediment and biota.

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While export to the basin from Arctic shelves is significant, up to 78% of the primary production and secondary production taking place on continental shelves is transferred to underlying shelf sediments (Fig. 5.3) (Berner, 1982; Walsh & McRoy, 1986; Hedges & Keil, 1995). Rates of organic matter transformation and nutrient regeneration in sediments play an important role in returning nutrients to the water column. Denitrification, the conversion of oxidized forms of inorganic nitrogen (nitrate, nitrite or nitrous oxide) to dinitrogen gas, is a major marine sink for combined nitrogen (Devol et al., 1997). Denitrification is enhanced in shelf and slope sediments so that the extensive Arctic shelves may contribute significantly to global rates of denitrification (Christensen et al., 1987). Ice is a key attribute linked to cycles of primary productivity on continental shelves. Marginal ice zones (MIZ) are some of the most dynamic areas in the world’s oceans with large seasonal and inter-annual fluctuations in ice cover and ice transport. These are unique habitats with specific communities dependent upon the physicochemical structure of the ice (Thomas & Dieckmann, 2002).

Fig. 5.3 Processes governing production, biogeochemical transformations, and fluxes of organic matter on the shelves and slopes of the Arctic Ocean. MIZ = Marginal ice zone. Modified from Anderson and Dyrssen (1989) and Grebmeier et al. (1998).

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MIZs in the northern Bering and Chukchi Seas and portions of the Barents Sea shelf are associated with some of the most productive ecosystems in the world (Grebmeier et al., 1995). The primary production associated with the MIZ, north of the Barents Sea polar front area, is typically between 40 and 50 g C m−2 y−1, while in the south of the polar front, primary production is typically 90–120 g C m−2 y−1 (Rey et al., 1987; Falk-Petersen et al., 1990; Wassmann & Slagstad, 1993; Hegseth, 1998). High bioproduction in the Barents Sea is due to several factors including (1) high annual primary production in close association with the receding ice edge and strong mixing following the breakdown of a previously stratified water column, (2) advection of large herbivorous zooplankton from the Norwegian Sea into the Barents Sea, and (3) transport of ice fauna by the Transpolar Drift from the Arctic Ocean into the Barents Sea where organisms are released during the melting process. This primary production consists of three components: (a) actively growing phytoplankton at the outer edge of the ice margin and in larger leads, (b) a layer of specialized sub-ice algal assemblage in pack ice, and (c) a sub-ice algal assemblage associated with multiyear ice (Syvertsen, 1991; Melnikov, 1997; Hegseth, 1998; Falk-Petersen et al., 2000b). The spring phytoplankton bloom in the Arctic results in a seasonal pulse of organic material that is short in duration but of high magnitude. Further north, a greater proportion of production takes place in ice rather than in water. In these areas, icealgae have been shown to act as an early-season carbon source that initiates biological production prior to significant phytodetritus flux (Ambrose & Renaud, 1997; Rowe et al., 1997; Stabeno & Overland, 2001). Studies from the Chukchi Sea (Ambrose & Clough, unpublished data) and Bering Sea (Stabeno & Overland, 2001) suggest that the duration of ice cover strongly affects the timing of the spring bloom, the predominant source of primary production (ice algae vs. phytoplankton) and the amount of local primary production exported to the benthos (Fig. 5.4). In heavy ice years, an early ice-associated spring bloom occurs and a larger fraction of the total production is due to ice algae rather than pelagic phytoplankton. Pelagic herbivores in turn respond to variations in the primary production regime. Large-bodied, energy-rich copepods such as Calanus glacialis and Calanus hyperboreus overwinter in deeper waters off the shelf break. Because they cannot complete their life cycles on the shelf itself, they must be advected onto shelves during the spring bloom. Their populations are favoured during blooms of large diatom-based phytoplankton, which tend to occur under high nutrient conditions. In contrast, primary production dominated by smallersized phytoplankton favours smaller, coastal zooplankton such as Calanus finmarchicus which have a far lower energetic value as a food source to higher trophic levels compared to the larger copepods. Pelagic arctic animals have adapted to strong seasonality by storing large amounts of lipids as energy reserves. The proportion of lipid to dry body mass

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Scenario 2: Abundant Ice

Scenario 1: Limited Ice

Phytoplankton

Primary Production

Ice Algae

Phytoplankton

Ice Algae

birds walrus fish

Zooplankton

Zooplankton seals, whales

Benthos

shrimp

Benthos

Fig. 5.4 Idealized scheme for the Norwegian Arctic describing two primary production scenarios (primary production dominated by phytoplankton vs. ice algae) related to ice cover, and their trophic implications. Thicker arrows and boxes signify primary pathways of energy (organic carbon) flows.

increases from 10 to 20% in phytoplankton to more than 50% in herbivorous zooplankton and ice-associated fauna. Lipids contained in copepods are then transferred through the food web, providing energetic needs for higher trophic levels. This lipid storage is one of the most fundamental specializations in polar bioproduction (Falk-Petersen et al., 2000a; Sargent & Falk-Petersen, 1988), and a major reason for the biodiversity of Arctic populations of zooplankton, fish, mammals and birds (Lavoie et al., 1999; Falk-Petersen et al., 1990). The supply of organic matter from surface waters also supports the energetic requirements of benthic communities in the Arctic. The benthos is highly dependent on the seasonal export of organic material provided by the spring bloom (Grebmeier & Barry, 1991; Wassmann et al., 1996) which in mid-Arctic locations tends to follow the retreating ice edge (Niebauer, 1991) or is associated with polynyas (Ambrose & Renaud, 1995; Piepenburg et al., 1995). Arctic benthic biomass is spatially variable and correlated with locations of high sediment carbon concentrations (Grebmeier & Barry, 1991; Wassmann et al., 1996). Grazing by herbivorous zooplankton greatly determines how various sources of primary production partition between pelagic and benthic systems (Mitchell et al., 1991; Sakshaug, 1991; Wassmann & Slagstad, 1993). Zooplankton grazers that are tightly coupled to phytoplankton blooms, as often seen in areas not influenced by ice, can drive a food web towards fish and bird production, whereas the absence of such grazers leads to benthic–pelagic coupling and a food web dominated by the benthos (Aagaard et al., 1999). Early-season herbivorous zooplankton and ice fauna are less abundant and do not graze icealgae efficiently, resulting in a large export of primary production to the benthic community. In contrast, in years with less ice, ice algae are of less importance, and open water phytoplankton blooms dominate the local primary

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production. The zooplankton community is matched to a later-occurring phytoplankton-based bloom, efficiently grazing such pelagic-dominated production and retaining much of the production in the water column, thus resulting in less carbon export to the bottom (Fig. 5.4). Direct physical disturbance events can also play a strong role in shaping benthic communities on local scales. Major disturbance from iceberg scour in the Arctic and Antarctic caused shifts in community structure from one of high diversity to one dominated by pioneer species, with effects lasting up to a decade (Gutt et al., 1996). Both direct disturbance and indirect effects (e.g. local temperature and salinity anomalies) have the potential to play a central role in shaping benthic community composition at smaller scales. 5.4

Environmental changes

Environmental conditions in the Arctic, as elsewhere, are not static. Indeed, environments change over a large range of ecologically relevant scales: diurnally, seasonally, annually, decadally, and over centuries. The patterns and scales of variability of the physical environment set the framework conditions of species existence and community interactions. Arctic organisms, communities and ecosystems are thus adapted to the patterns of their respective local climates and the inherent environmental variability. Under such conditions, natural systems can be in quasi-equilibrium with the climate forcing mechanisms. In contrast, environmental changes greater than the range of adaptability of Arctic systems influence patterns of biodiversity directly via induced mortality (Harper, 1977; Sousa, 1979; Dethier, 1984; Sousa, 1984; Gutt et al., 1996), and indirectly through a complex of feedback loops altering interactions between existing species (Bond et al., 1992; Menge, 1995; Carroll & Highsmith, 1996). Such environmental changes can be either via extreme events such as storms and severe temperature anomalies, or long-term environmental changes such as global climate change and habitat modification. Further, effects of environmental changes may be exacerbated in species living in sub-optimal habitats at or near their limits of physiological tolerance (Bowman & Lewis, 1977; Lewis et al., 1982; Sousa, 1984; Woodward, 1987; Carroll, 1994), as is commonly the case in the Arctic marginal seas (Zenkevich, 1927; Dahle et al., 1998; Smith et al., 1999). 5.4.1

Climate variability

Several types of Arctic climate variability are known to exist. We define variability herein as cycling or oscillation between relatively stable states. Examples range from rapid changes and regime shifts to oscillations extending from high to low frequency and to long-term trends (Dickson et al., 1988; Ebbesmeyer et al., 1990). The existence of inter-annual variations is well documented in the meteorological (Dement’ev, 1991) and oceanographic literature

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(Treshnikov & Baranov, 1976; Yanes, 1977; Nikiforov et al., 1989; Midttun & Loeng, 1991; Parkinson, 1991). Considering that ecological systems are generally non-linear (May, 1986), even modest changes in large-scale climate variability may produce large effects at various trophic levels (Ottersen et al., 2001). Interannual and decadal timescale variations in the Arctic are under the primary influence of three dominant climatic forcing mechanisms: the Arctic oscillation (AO) (Thompson & Wallace, 1998), North Atlantic oscillation (NAO) (Hurrell, 1995) and the Pacific decadal oscillation (PDO) (Mantua et al., 1997). The AO is related to, and defined by, the wintertime sea level pressure field centred over the Arctic. This pressure field can flip back and forth several times between high and low pressure each winter. Low pressure is associated with westerly winds in the upper atmosphere, keeping colder air in the far north and resulting in the encroachment of warmer weather northward. In contrast, high pressure over the central Arctic enables cold air to seep farther southward into lower latitudes (i.e. continental Europe). Similarly, Proshutinsky and Johnson (1997) and Johnson et al. (1999) have identified a decadal variation in Arctic ice and ocean circulation, with wind-driven motion in the central Arctic alternating between an anti-cyclonic circulation regime (ACCR) and cyclonic circulation regime (CCR), with each persisting for 5–7 years. Shifts from one regime to another are forced by changes in the location and intensity of the Icelandic low, Siberian high and Aleutian low pressure systems. These transformations from one regime to another can be defined as climate shifts in the Arctic Ocean, and occur quite rapidly. The anti-cyclonic regime is characterized by a cold and dry Arctic atmosphere and cold and saltier polar ocean, whereas the cyclonic regime is characterized by a warm and wet atmosphere and warm and fresh polar ocean. The AO obviously interacts with the NAO and PDO at the margins of the Arctic (Dickson et al., 2000). At inter-decadal scales (15–20 years), Mysak et al. (1990) and Mysak (1995) linked Arctic variations to those in the North Atlantic, while (Serreze et al., 2000) noted an 81% correlation between AO and NAO indices, suggesting ‘existing evidence indicates that the NAO can be considered as a major component of the AO’. Marine ecosystems respond directly and indirectly to climate variability acting through modification of ocean dynamics (Smith et al., 1999), and there is no doubt that climatic variations affect biodiversity and species distributions. The distribution of the cod stock in western Greenland expanded northward with warming ocean temperatures in the 1920s and 1930s (Jensen, 1939). After a 30-year period of steady ocean temperatures, cooling in the Labrador Sea beginning in the 1960s has severely reduced the cod stock along their former range (Buch et al., 1994). Norwegian herring stocks had a feeding migration to Iceland during this warm period in the 1920s through the 1950s, but these stocks gradually disappeared from Iceland following a subsequent climate shift which resulted in lower temperatures in the 1960s (Vilhjalmsson, 1997). There are also a number of

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examples of biogeographic shifts of Arctic benthic communities with local species compositions favouring Atlantic or Arctic organisms correlating with periods of warmer or colder water respectively (Blacker, 1957; Galkin, 1964, 1998; Dyer et al., 1984). In the Bering Sea, we have the best-documented examples of climatic regime shifts cascading through all trophic levels of the marine ecosystem (Hare & Mantua, 2000). In 1977, the PDO abruptly shifted from a cold to a warm phase, which raised the sea surface temperature and suppressed primary production. This in turn reduced the abundance of lower trophic level species such as shrimp, crab and herbivorous fish such as capelin. The reduced numbers of lower trophic-level prey may have led to declines in the numbers of Steller’s sea lions and harbour seals. The declining numbers of sea lions and seals apparently led to increased predation on sea otters by killer whales, resulting in a drastic reduction (>70%) in the population of Alaskan sea otters. The sea otter population crash allowed sea urchins (a major part of the sea otter diet) to proliferate. Sea urchins feed on kelp; so the kelp forests in the Gulf of Alaska and the Bering Sea were overgrazed, which affected all of the organisms associated with the kelp including seabirds like puffins and kittiwakes. In 1997, combined effects of the AO and El Niño–Southern Oscillation (ENSO), resulted in relatively calm winds and reduced cloud cover over the Bering Sea. The increased sunlight caused the surface water temperature to rise more than 2°C above normal, but few nutrients were present in the stratified water. These conditions favoured a coccolithophore bloom instead of diatoms which typically form the base of the summer food web. Since coccolithophores became established in the Bering Sea in 1997, they have been present in subsequent years even though weather conditions have returned to more normal states. A major ecosystem shift has also occurred in the Bering Sea in the past two decades in response to warmer sea temperatures (Stabeno & Overland, 2001). Less sea ice is present and an earlier winter-to-spring has resulted in long-term changes in fish populations (Hare & Mantua, 2000) with concomitant impacts on the commercial fishing industry. 5.4.2

Long-term climate change

The interaction between the Arctic Ocean and the global climate system has long been recognized. Fully a century ago in 1902, Fridtjof Nansen wrote: ‘It is evident that oceanographic conditions of the North Polar Basin have much influence on climate, and it is equally evident that changes in the conditions of circulation would greatly change the climatic conditions’. In the ensuing century, global mean temperatures have risen by about 0.6°C with more than half of the increase occurring in the past 25 years (Serreze et al., 2000). Significant climatic changes in the Arctic have been observed over the past 35 years, although the trends are regionally variable: areas such as Alaska

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and western Russia have warmed by more than +1°C per decade, while others, such as eastern Canada, Southwest Greenland and the Labrador Sea have exhibited cooling trends (Chapman & Walsh, 1993; Serreze et al., 2000). Climate models predict that global climatic changes will be mostly pronounced in polar regions (IPCC, 1998). It is thought that the impact of warming will be amplified in the Arctic due to the combined effects of sea ice retreat and stable atmospheric stratification (Manabe & Stouffer, 1994; Rind et al., 1995; Weller & Lange, 1999). Several lines of evidence indicate that significant environmental change is already occurring in terrestrial and marine Arctic and sub-Arctic environments (Grebmeier & Dunton, 2000; Morison et al., 2000; Serreze et al., 2000). Decreased sea-ice extent, ice thickness, salinity, and increased water and air temperature have been documented over the past two decades, as have shifts in water mass boundaries, ocean circulation and atmospheric circulation (Carmack et al., 1995; McLaughlin et al., 1996; Cavalieri et al., 1997; Overpeck et al., 1997; McPhee et al., 1998; Melling, 1998; Rothrock et al., 1999). These patterns of arctic change generally agree with those predicted by current climate models under scenarios of enhanced greenhouse warming. Because the Arctic is a key element in global climate regulation, small perturbations to the present day climate system may be dramatically amplified in the Arctic. The return flow of North Atlantic water northward via the Gulf Stream is one of the primary determinants of climate in continental Europe. Global circulation models for the eastern Arctic show that the inflow of warm water via the North Atlantic is extremely sensitive to changes in salinity and temperature (Broecker, 1990, 1991, 1994; Manabe & Stouffer, 1995). The large-scale currents driven by thermohaline circulation in the North Atlantic may significantly weaken or abruptly shut off completely in response to the byproducts of global warming (Broecker, 1994, 1997), resulting in a significantly colder Barents Sea region. As much of the freshwater in the Arctic Ocean is stored in sea ice, changes in the heat budget of the ice cap are expected to lead to changes in the rate of freshwater export, and hence, nutrient and organic matter exports to the basin (Walsh et al., 1997). The annual primary production is coupled to the spatial variation in ice cover, and so will vary between warm and cold years (Slagstad & Støle-Hansen, 1991; Slagstad & Stokke, 1994; Sakshaug & Walsh, 2000). The trend towards a decreased extent and thickness of the sea ice cover in the Arctic (Rothrock et al., 1999; Wadhams & Davis, 2000) poses serious threats to ice-related organisms dependent upon the physicochemical environment provided by sea ice. Additionally, the effects on higher trophic levels of the Arctic ecosystems could be dramatic if the partitioning of the primary production between ice-related and pelagic components changes. Less ice may lead to a less pronounced melt-water layer, and possibly earlier melting (Stabeno & Overland, 2001). This will influence the timing and duration of the pelagic

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spring bloom (Engelsen et al., 2002) and the match–mismatch between primary and secondary production (Slagstad & Støle-Hansen, 1991; Wassmann & Slagstad, 1993; Falk-Petersen et al., 1999). Because marine ecosystems in the Arctic have lower species diversity and are composed of relatively simple food webs with a limited number of trophic levels (Fig. 5.5), they are considered

(4,5) Top Carnivore Polar Bear

(3,4) Ringed Seal (3) Benthos Feeding Marine Mammals Walrus Bearded Seal Gray Whale (3) Starfish (3) Benthos Feeding Marine Birds Eiders

(3) Shorebirds (2) Benthos Infauna: Annelids Bivalves Gastropods Epifauna: Mysids Amphipods Isopods Crabs Shrimp Echinoderms

(4,5) Predatory Birds Gulls

(4) Piscivorous Marine Mammals Spotted Seal Beluga Whale Narwhal Harp Seal Harbor Seal N. Sea Lion Dall′s Porpoise

(3) Fish & Squid Marine: Cods Pollock Herring Capelin Sandlance Flounders Sculpins Squid Anadromous: Salmon Arctic Char Ciscos Smelt Whitefish

(3,4) Marine Birds Murres Kittiwakes Cormorants Puffins

(3) Planktivorous Marine Mammals Bowhead Whale

(2) Ice Invertebrates Amphipods Isopods Mysids

(3) Planktivorous Marine Birds Auklets

(2) Zooplankton Copepods Euphausiids Crab & Shrimp larvae

(1) Phytoplankton, Ice Algae, Macrophytes, Detritus

Fig. 5.5 Generalized food web of the Arctic marine ecosystem, showing major categories of biota and specific examples. Specific organisms vary with region and season. Numbers in parentheses indicate trophic levels in ascending order. Modified after Becker (1994).

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less stable to perturbations than corresponding ecosystems with higher diversity (Johnson et al., 1996). Thus, modifications to the primary production regime can ripple throughout the food web. 5.4.3

Ozone and ultraviolet radiation

Data available indicate that ultraviolet-B radiation (UVB) levels have increased significantly over the past 10–15 years at mid-latitude areas of the northern and southern hemispheres (Crutzen, 1992; Kerr & McElroy, 1993; Madronich et al., 1995; Wardle et al., 1997). These increases in UVB are linked to reductions in stratospheric ozone (Kerr & McElroy, 1993; Madronich et al., 1995; Wardle et al., 1997). Severe seasonal reductions in ozone layer thickness have been recorded over the Arctic (Fergusson & Wardle, 1998; Goutail et al., 1999). A growing number of studies indicate that UVB radiation at current levels is harmful to aquatic organisms and may reduce the productivity of marine ecosystems (Holm-Hansen et al., 1993; Siebeck et al., 1994; Häder et al., 1995). Such UVB-induced decreases in productivity have been reported for phytoplankton, heterotrophs and zooplankton, the key intermediary levels of marine food chains (Damkaer, 1982; Thomson, 1986; Smith & Cullen, 1995; Booth et al., 1997; Häder, 1997; Browman et al., 2000). UVB-induced changes in food chain interactions can be far more significant than direct effects on individual organisms at any single trophic level (Bothwell et al., 1994; Hessen et al., 1997; Williamson et al., 1999). UVB exposure reduces the total lipid content of some microalgae (Arts & Rai, 1997; Plante & Arts, 1998), and this effect includes the polyunsaturated fatty acids (PUFAs) (Goes et al., 1994; Wang & Chai, 1994; Hessen et al., 1997). These PUFAs cannot be synthesized by metazoans, and appear highly important for the entire marine food web including zooplankton and fish (Rainuzzo et al., 1997; Sargent et al., 1997). 5.4.4

Contaminants

Although considered relatively pristine by most standards, contaminants found in remote Arctic areas include persistent organic contaminants, heavy metals (lead, cadmium, mercury), acidifying gases and radionuclides (AMAP, 1997; Macdonald et al., 2000). The Arctic ecosystem’s unique features may lead to increased vulnerability from contaminants. These include the seasonal and spatial focus of primary productivity, relatively simple food web structure, strong benthic–pelagic coupling, a prevalence of large mammals as top predators and relatively high lipid contents in some species. Long-range transport of contaminants from hemispheric to global scales has been identified as a key process leading to increased concentrations of anthropogenic hazardous chemicals in the northern polar region. Low temperatures in the Arctic seem to create a sink for certain persistent organic pollutants, which may in some cases result in contaminant levels that are higher in the Arctic than in the source

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regions (cold-condensation effect) (Wania & Mackay, 1996). Biological, chemical and ecological factors influence to varying degrees, depending on species, the trophic transfer of contaminants through Arctic marine food webs (Dietz et al., 2000; Fisk et al., 2001; Borga et al., 2002). In relation to global change, alterations in species energetics and trophic interactions may affect the bioamplification and delivery of pollutants to consumers of arctic fish, mammals and bird populations. 5.5

Natural resources and ecological services

Identifying and valuing natural resources and ecosystem services to mankind is a vast and complicated blend of economics and environmental science generally termed ecological economics. Ecological economists have drawn attention to the services and their economic value provided by global ecosystems. These ecological services have been described as natural capital. According to Costanza et al. (1997), the total value of ocean and coastal ecosystem services is 21 trillion US dollars, representing 60% of the estimated value of the global ecosystem services. The Arctic Ocean and its marginal seas provide a number of unique ecological services as part of the global life support system connected to the oceans. Here, we shall try to tie the natural science together in terms of the ecological services they provide to the local and global populations. According to Knapp (2000), the total population of northern regions in 1990 was 10.5 million inhabitants. The vast majority of this northern population (85%) is found in Russia. Of the total population, 1.1 million, or about 11%, is defined as indigenous. The proportion of indigenous peoples to total population varies from a high of 83% in Greenland to a low of 5% in Scandinavia. 5.5.1

Indigenous people

Indigenous people of the Arctic, through culture and history spanning, perhaps 1000 years or more, directly and indirectly rely on arctic natural resources and ecological services more heavily than other population groups. Although individual native populations in the Arctic are extremely diverse in their culture and history, subsistence-based harvesting of natural resources has been, and continues to be, an essential part of their lifestyle. In the past, people were totally dependent upon subsistence harvesting of natural resources for survival in the Arctic. Because of this dependency, the practical and spiritual connection to the land, sea, and its resources and environment was immense. Today, although most native populations are not totally dependent upon subsistence harvesting for survival, subsistence is still a central part of the fabric of society (Watson, 1997; Caufield, 2000) and essential for economic survival, social identity and spiritual life (Nuttall, 2000). Although store-bought goods

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are now available in most native towns, consumption of locally harvested resources still accounts for a large proportion of the diet of many native households. Harvested resources also can provide cash income to households and goods to barter in the mixed subsistence-cash economies prevalent today (Young, 1992; Alaska Natives Commission, 1994). Subsistence resources and the activities related with the harvest of these resources are also associated with spiritual values and community well-being of northern people (Callaway, 1999). Obviously, for a group so closely tied to the local resources, any environmental changes that affect the abundance, distribution and timing of availability of resources have strong implications to the culture. Four main concerns related to environmental change in the marine realm, recently identified by Alaskan native populations, are probably accurate for the Arctic as a whole: (1) changing patterns of fish stocks and other harvestable resources such as marine mammals, (2) the distribution of sea ice, which marine mammals utilize as haul outs and hunters use in pursuit of prey, (3) pollution and contaminants, and (4) frequency and strength of storms, which can put coastal villages at higher risk. 5.5.2

Non-indigenous regional populations

Residents of the Arctic, while not dependent upon subsistence harvesting of resources in the manner of native groups, still rely on harvesting of natural resources. Perhaps one of the main ecological services of the Arctic to local residents is through the stocks of commercially harvested fish. Although not significant on a global scale according to the FAO, commercial fishing in the northern Arctic waters has tremendous importance to the wellbeing of northern populations. Norway and Russia share one of the most prolific of global cod harvests in the Barents Sea, with a current harvest of about 400 000 tons annually (Institute of Marine Research Bergen, 2001). Other fish stocks such as capelin and herring also contribute immensely to the local economic basis of the population surrounding the Barents Sea. Greenland is the world’s largest exporter of shrimp, and commercial fishing and fish processing is the most important industry in the country, with a per capita economic value of US$4200 (Grønlands Statistik, 1999). In Alaska, the Bering Sea commercial fishing industry is integral to the local economy. The seafood industry accounts for 11% of Alaska’s total employment, more than any other private sector activity, and has a total value (landing and processing) of almost $3 billion. Fish landings in Alaska account for 54% of total volume of US landings and 37% of the total ex-vessel value of total US landings (Knapp et al., 1999). Marine fisheries in the Arctic are vulnerable to climate change. Marine fish harvests are almost totally dependent on climate-related environmental

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conditions affecting fish abundance. Significant and strong correlations have been found between the abundance of various fish stocks and ocean temperatures, air temperatures, location and intensity of major air masses such as the Aleutian low in the Pacific or the NAO in the North Atlantic, and sea ice distribution. The mechanisms, however, between climate and fisheries abundance are not well understood and probably vary by species and region. They may include changes in the velocity and direction of ocean currents, affecting the availability of nutrients and the fate of larval and juvenile organisms, changes in ocean temperatures that affect the abundance of both harvested species and their predators and prey, and changes in UVB radiation that may have negative effects on fish eggs that float near the surface. Even the most responsibly managed fisheries exhibit substantial climate-related variations in harvests over time. Climate change may reduce the abundance of some species, while simultaneously enhancing different stocks. This strongly challenges fisheries management paradigms and principles in order to take account of climate change effects when making fisheries management decisions. 5.5.3

National/international/global users

The Arctic is also important to the global community. Oil and gas extraction is a significant industrial activity in the Arctic and supplies the world with petroleum-based products. Russia, the United States, Canada, Norway and Greenland are either actively producing or are exploring for oil or gas in their Arctic areas. A large portion of known and expected petroleum reserves in the Arctic is in marine areas. Currently, environmental conditions such as sea ice are, at best, a major engineering challenge in the extraction of petroleum from the Arctic seas, and at worst result in an inability to carry out activities in certain areas. If ice conditions become less extensive as a result of changing environmental conditions in the Arctic, this may allow more extensive activity than in the past. However, the marine ecosystems may also change in the same areas, possibly resulting in greater conflicts between industrial activity and biological resources, and on the industries that rely on them (i.e. fishing). Ice conditions presently prohibit widespread commercial use of the Arctic Ocean as a commercial shipping route. While limited internal traffic exists to resupply remote villages during the summer, virtually no commercial freighter traffic transits the northeast passage north of Russia as a through route from the North Atlantic to Asia. An icebreaker must accompany any freighters using this route. If ice conditions become less severe, it would be possible to significantly increase the flow of commercial traffic through the northeast passage, saving an average of 13 days on the journey between northern Europe and Asia (Kitagawa, 1996). A similar possibility exists on the northwest passage through the Canadian Archipelago. Shipping also has negative potential consequences, via the increased risk of pollution or a shipping

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accident (Norwegian Maritime Directorate, 2000), or through possible introduction of alien species through hull fouling or ballast water exchanges (Lavoie et al., 1999; Wonham et al., 2001). Acknowledgements The authors wish to acknowledge Vladimir Savinov for preparing graphics and figures and Barbara Vögele for graphics support and in managing the reference material and literature citations. References Aagaard, K. and Carmack, E.C. (1989) The role of sea ice and other fresh water in the Arctic circulation. Journal of Geophysical Research, 94, 14485–14498. Aagaard, K., Coachman, L.K. and Carmack, E.C. (1981) On the halocline of the Arctic Ocean. Deep-Sea Research, 28, 529–545. Aagaard, K., Darby, D., Faulkner, K., Flato, G., Grebmeier, J., Measures, C. and Walsh, J. (1999) Marine Science in the Arctic: a strategy, Arctic Research Consortium of the United States (ARCUS), Fairbanks, AK. Alaska Natives Commission (1994) Alaska Natives Commission Final Report, 3 volumes. AMAP (1997) Arctic Pollution Issues: a state of the Arctic environment report, Arctic Monitoring and Assessment Program, Oslo, Norway. Ambrose Jr., W.G. and Renaud, P.E. (1995) Benthic response to water column productivity patterns: evidence for benthic-pelagic coupling in the Northeast Water Polynya. Journal of Geophysical Research-Oceans, 100, 4411–4421. Ambrose Jr., W.G. and Renaud, P.E. (1997) Does a pulsed food supply to the benthos affect polychaete recruitment patterns in the Northeast Water Polynya? Journal of Marine Systems, 10, 483–495. Anderson, L. and Dyrssen, D. (1981) Chemical constituents of the Arctic Ocean in the Svalbard area. Oceanologica Acta, 4, 305–311. Anderson, L. and Dyrssen, D. (1989) Chemical oceanography of the Arctic Ocean, in The Arctic Seas (ed. Y. Herman), Van Nostrand Reinhold, pp. 93–114. Anderson, L.G. (1995) Chemical oceanography of the Arctic and its shelf seas, in Arctic Oceanography: Marginal Ice Zones and Continental Shelves (eds W.O. Smith Jr. and J.M. Grebmeier), American Geophysical Union, Washington, DC, pp. 183–202. Anderson, L.G., Olsson, K. and Chierici, M. (1998) A carbon budget for the Arctic Ocean. Global Biogeochemical Cycles, 12, 455–465. Anderson, R.F., Lyons, T.W. and Cowie, G.L. (1994) Sedimentary record of a shoaling of the oxic/ anoxic interface in the Black Sea. Marine Geology, 116, 373–384. Arts, M.T. and Rai, H. (1997) Effects of enhanced ultraviolet-B radiation on the production of lipid, polysaccharide and protein in three freshwater algal species. Freshwater Biology, 38, 597–610. Becker, P.R. (1994) Characterization of the Arctic environment. Marine biological resources. Arctic Research of the United States, 8, 66–76. Berner, R.A. (1982) Burial of organic carbon and pyrite sulfur in the modern ocean: its geochemical significance. American Journal of Science, 282, 451–457. Björk, G. (1989) A one-dimensional time-dependent model of the vertical stratification of the upper Arctic Ocean. Journal of Physical Oceanography, 19, 52–67.

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6

The Arabian Sea S.W.A. Naqvi, Hema Naik and P.V. Narvekar

6.1

Introduction

The Arabian Sea makes contributions to global biogeochemical cycles that are far in excess of what might be expected from its modest size. It experiences extremes in atmospheric forcing, asymmetrically distributed over the region, that bring about exceptionally large hydrographical changes and produce a wide variety of ecosystems or biogeochemical provinces. The diversity and spatio-temporal proximity of these provinces make the region a natural laboratory to investigate present biogeochemical processes, and to apply this knowledge for reconstructing past changes, as well as for predicting future responses of oceanic ecosystems to human-induced climatic change. It has therefore attracted a great deal of attention of oceanographers world over, and has been subjected recently to numerous investigations under the aegis of the joint global ocean flux study (JGOFS). The JGOFS results supplemented by those of two other contemporary programmes – the global ocean ecosystem dynamics (GLOBEC) and the world ocean circulation experiment (WOCE) – have immensely improved our understanding of biogeochemistry of this region. Relying largely on these results, this chapter focuses on selected topics of general interest. For detailed accounts, the following seminal literature is recommended: Zietzschel (1973); NIO (1977); Angel (1984); Huq and Milliman (1984); Smith (1998, 1999, 2000, 2001); Desai (1992); Ittekkot and Nair (1993); Burkill etal. (1993); Baars (1994); Lal (1994); Van Veering etal. (1997); Krishnaswami and Nair (1996); Burkill (1999); Gaur (2000); Gage et al. (2000); Pfannkuche and Lochte (2000); Sen Gupta and Desa (2001); Watts etal. (2002). 6.2

Geographical setting

Situated in the northwestern Indian Ocean, the Arabian Sea is bounded by the African and Asian landmasses to the west and north and by the Indian subcontinent to the east (Fig. 6.1). Unlike these natural boundaries, the southern boundary that separates the Arabian Sea from the greater Indian Ocean is arbitrarily defined. For the oceanographic purpose it is generally taken to run from Goa (India) along the western side of the Laccadive and Maldive Islands to the equator, and thence slightly to the south to Mombassa (Kenya) (Schott, 1935). The region so demarcated occupies an area of 6.2 × 106 km2. It does not include the Gulfs of Aden and Oman through which the Arabian Sea is

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RHJ 20° N

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Fig. 6.1 Major features of surface circulation in the NW Indian Ocean during (a) Northeast monsoon, and (b) Southwest monsoon (EACC – East African Coast Current; SECC – South Equatorial Countercurrent; SC – Somali Current; WICC – West Indian Coast Current; NMC – Northeast Monsoon Current; LH – Lakshadweep High; SG – Southern Gyre; GW – Great Whirl; SE – Socotra Eddy; RHJ – Ras-al-Hadd Jet; LL – Lakshadweep Low; SMC – Southwest Monsoon Current). Redrawn from Schott and McCreary (2001) with permission from Elsevier Science.

connected to two Mediterranean-type marginal seas – the Red Sea and the Persian Gulf. Also excluded by the above demarcation is the Laccadive Sea, a smaller water body (area 0.23 ×106 km2) that lies to the east of the Laccadive Islands (Lakshadweep). However, oceanographers generally do not make a distinction between the Arabian Sea and the Laccadive Sea especially while dealing with the processes along and off the continuous west coast of India; such a distinction will not be made here. The most prominent bathymetric feature of the Arabian Sea is the northwest– southeast trending Carlsberg Ridge that divides the Arabian Sea into two major (depth> 4000 m) basins – the Arabian Basin in the northeast and the Somali Basin in the southwest (Fig. 6.1). While almost the entire Arabian Basin is located within the Arabian Sea, a large portion of the Somali Basin falls outside. A less pronounced feature is the Murray Ridge that extends southwest from the Makran margin to join the Carlsberg Ridge thereby separating the Arabian Basin from the relatively narrow and shallow (depth < 2000m) Oman Basin. The continental shelf is generally wide (often exceeding 100km), east of Karachi along the Pakistani coast and all along the Indian west coast with the maximum (350km) occurring off the Gulf of Cambay. Elsewhere, the shelf width rarely exceeds 40km. 6.3

Climate and circulation

The landmass that limits the poleward expanse of the northern Indian Ocean to just about 25° N latitude profoundly affects physical processes in the Arabian Sea.

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The continental influence is best manifested by the monsoons, the seasonally reversing wind systems that are driven by pairs of low (high) and high (low) pressure zones over the sea and land, respectively, during winter (summer). These originate as a result of the more rapid cooling (heating) of the land relative to the sea during winter (summer). Of the two monsoons, the summer or southwest monsoon (SWM) is far more energetic. During this period (June–September), strong winds blow from the southwest with speeds frequently exceeding 30knots, especially along a strongly sheared low-level atmospheric jet (the Somali Jet – Findlater, 1971), the axis of which extends from the Somali coast towards the Gulf of Cambay. In comparison, the winter or northeast monsoon (NEM) is characterized by weaker winds that blow from the northeast during December–February. The periods intervening the monsoons – March–May and October–November – are referred to as the spring intermonsoon (SI) and fall intermonsoon (FI) respectively; during these periods the winds are light, and slowly reverse direction. Almost all the rainfall over the Arabian Sea occurs during the SWM and along its eastern shores where it may exceed 300 cm y−1. The amount of precipitation decreases towards the northwest, and so the balance between evaporation and precipitation is at its maximum off the Arabian coast and at its minimum along the southwest coast of India. The Arabian Sea does not receive much river runoff, the combined discharges by the main rivers (the Indus, the Narmada and the Tapi), all draining into the northeastern Arabian Sea, probably not exceeding 200 km3 y−1. However, there are scores of other small rivers originating in the Western Ghats (a mountain range separating the narrow western coastal plain from the Deccan Plateau), which together transport about 150 km3 of freshwater annually (most of it during the SWM period). The large rainfall and land runoff combine to result in a positive water balance (excess of precipitation and runoff over evaporation) over a few hundred kilometre wide belt along the Indian coast. The net water balance is negative elsewhere, and for the Arabian Sea as a whole. Consequently, the surface waters are the least saline in the southeast and the most saline in the northwest (Wyrtki, 1971). Associated with seasonal changes in the wind field, the near-surface oceanic circulation also reverses completely every six months. The tropical nature of the North Indian Ocean and its relatively small size affect the dynamics of surface waters in another important way. Changes in the monsoon winds generate coastal and equatorial Kelvin waves and equatorial Rossby waves, having both annual and sub-annual periods, which propagate rapidly through the region, strongly influencing circulation at sites far away from their origins. The combination of such remote forcing and local wind forcing produces features of surface circulation not observed in other oceanic areas. This, in turn, has a profound impact on biogeochemical processes. Major surface currents in the Arabian Sea during the two monsoon seasons are schematically shown in Fig. 6.1 (Schott & McCreary, 2001). During the

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SWM, near-surface circulation in the region is generally clockwise. The most energetic flow during this season occurs off Somalia with a volume transport comparable to that of the Gulf Stream. The Somali current system includes two major, quasi-stationary anti-cyclonic eddies, the southern gyre (SG) and the great whirl (GW) [a third, the Socotra Eddy (SE), may also be found in the vicinity of the island it has been named after]. These eddies cause most of the northward coastal flow to be deflected to the east and south at two locations – around 4° and 10° N. The deflected flows feed the southwest monsoon current (SMC), a feature unique to the North Indian Ocean (it replaces the westward flowing north equatorial current found in the other oceans). Coastal upwelling associated with this flow regime is manifested by wedge-shaped features [discernible in satellite imageries of sea surface temperature (SST)] along the left (shoreward) shoulders of the SG and GW (Fig. 6.2f). As compared to 4° N, upwelling at 10° N is more pronounced and longer lasting. Further north, strong SWM winds force even more widespread upwelling along the coasts of Yemen and Oman (Figs 6.2f and 6.3c), but surprisingly the near-surface currents here do not exhibit an organized pattern and seasonality expected from the wind stress, except near the coast (e.g. the Ras-al-Hadd Jet off Oman). Instead, the flow seems to be dominated by eddies, having spatial scale of 300–400 km, which account for the bulk of the kinetic energy (Flagg & Kim, 1998). The presence of these eddies ensures that the cold upwelled water is rapidly transported offshore as filaments and plumes which extend several hundred kilometres from the coast (Figs 6.2f and 6.3c). The large-scale dispersal of coastally upwelled water, together with possible offshore upwelling forced by strong gradients in wind speed across the Somali Jet (Bauer et al., 1991), and entrainment of water from the thermocline due to turbulence, lowers SST over a very large area as evident from satellite imageries (Fig. 6.2f). A similar offshore spreading of upwelled water is not seen off the southwest Indian coast even though during the SWM this region is known to possess all characteristics of an eastern-boundary environment: coastal upwelling, equatorward surface flow [southward flowing West India coastal current (WICC)] and a poleward undercurrent (Shetye et al., 1990). The shoaling of thermocline occurs over a large area along the Indian coast and in the region of the Lakshadweep low (LL), a seasonal cyclonic eddy. But the cold, saline upwelled water is overlain by a thin (~10 m), warm, low-salinity layer formed as a result of precipitation and land runoff, because of which the signatures of upwelling are barely observable in the remotely sensed data. Moreover, the period of upwelling (May–November) extends well beyond the SWM season indicating that the process cannot be forced entirely by local winds. Model simulations suggest that remote forcing from the Bay of Bengal plays a major role in driving the WICC (McCreary et al., 1993). During the NEM, when the near-surface circulation is generally counterclockwise, the SMC is replaced by the westbound northeast monsoon current (NMC). Surface currents in the northern and western Arabian Sea are weaker

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Fig. 6.2 Monthly climatologies of remotely sensed surface chlorophyll a (a–c) and SST (d–f) for February (a, d), May (b, e) and August (c, f). Chlorophyll and SST climatologies were created by Jerry Wiggert (University of Maryland, College Park) and Bob Evans (RSMAS, University of Miami) using the SeaWiFS (Sea-viewing Wide Field-of-view Sensor) and MODIS (Moderate-resolution Imaging Spectroradiometer) data, respectively. Note that some features of circulation (e.g. the great whirl) can be readily identified in both sets of data (see Color Plate 2).

and less organized as compared to the SWM. However, in spite of much weaker winds, circulation in the eastern Arabian Sea is best developed during the NEM. The WICC, which flows northward into the wind, is obviously remotely forced. With a volume transport approximately twice that of the

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SWM, but in the opposite direction, the WICC during the NEM is broadest off the SW coast of India, where the LL is replaced by the Lakshadweep high (LH), gradually narrowing towards the north (Shetye et al., 1991). The warm, low-salinity waters of the WICC exert a major control on biogeochemical cycling. Outside the region affected by the WICC, the surface waters are cooled to 24–26°C (Fig. 6.2d) by the northeasterly continental winds, leading to deep mixed layers (MLDs) (Fig. 6.3a) and entrainment of water from the

NO3– (μM)

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Distance from Oman coast (km) Fig. 6.3 Distributions of temperature (°C, a–c) and nitrate (μM, d–f) in the upper 200 m along the US JGOFS southern transect (see inset of Fig. 6.4 for station locations) during the NEM (TN043; a, d), SI (TN045; b, e) and SWM (TN050; c, f).

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thermocline, a process prevented by strong thermohaline stratification in the WICC. Like the near-surface circulation, the subsurface water movement in the northern Indian Ocean is also greatly impacted by the existence of the low latitude northern boundary. In the absence of sites of deep convection north of the equator, the intermediate and deep layers in the region are ventilated mostly from the south. Although the Arabian Sea receives dense outflows from the Persian Gulf and Red Sea, their combined magnitude is no more than 0.64 × 106 m3 s−1, an order of magnitude lower than the value for intermediate waters derived from the southern hemisphere (Warren, 1994). These waters enter the Arabian Sea mostly along its western boundary and during the SWM (Swallow, 1984). 6.4 6.4.1

Nutrients and primary production Subsurface nutrient trap

The mesopelagic zone of the Arabian Sea is renewed at a surprisingly rapid rate (on a time scale of 1–10 years – Naqvi, 1987; Olson et al., 1993; Howell et al., 1997). However, the vigorous ventilation does not bring about sufficient re-oxygenation because the waters responsible for renewal lose most of their O2 content along their tortuous trajectories of flow from their formation sites and in the equatorial region before they enter the Arabian Sea (Swallow, 1984; Warren, 1994). The residual O2 is consumed rapidly for the degradation of copious amounts of organic matter produced within the Arabian Sea itself. Consequently, O2 levels fall very close to zero while nutrients accumulate in high concentrations within a zone that extends from the base of the surface mixed layer to a depth of approximately 1 km in the northern Arabian Sea. Due to the extremely steep gradients of nutrients in the upper thermocline [e.g. up to 1 μM nitrate (NO3− ) m−1; Fig. 6.3], minor perturbations can inject them in large quantities into the euphotic zone. The unique monsoon forcing provides the energy needed for this purpose. For example, summer upwelling raises surface NO3− concentrations up to 16–20 μM off Somalia (Smith & Codispoti, 1980), India (Naqvi et al., 1998) and Oman (Fig. 6.3f) (Morrison et al., 1998), whereas winter convection results in an enrichment of 4–6 μM NO3− in surface waters of the northern Arabian Sea (Fig. 6.3d) (Madhupratap et al., 1996b; Morrison et al., 1998; Naqvi et al., 2002). This greatly stimulates the growth of phytoplankton, leading to the formation of blooms, as evident from elevated chlorophyll (chl) levels derived from the remotely sensed ocean colour data both during summer and winter (Figs 6.2a and 6.2c). These blooms spread over a large area corresponding remarkably well to the zones of low SST (Figs 6.2d and 6.2f). The ensuing sedimentation of particulate organic carbon (POC) and its decomposition contribute to the maintenance of high nutrient

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concentrations in subsurface waters. The region thus serves as a very efficient nutrient trap where the low O2-high nutrient conditions in subsurface waters are tightly coupled to high surface productivity (Ryther et al., 1966). 6.4.2

Primary productivity

25

250

Latitude (N)

Primary productivity (mmole C m–2 d–1)

The elevated phytoplankton biomass inferred from satellite observations is corroborated by field measurements of primary productivity (PP) and chl a (Marra, 2002 and references therein). That the Arabian Sea is a region of high biological productivity has long been known (Ryther et al., 1966; Qasim, 1977, 1982). However, until recently high PP was believed to be fuelled mainly by nutrients supplied to euphotic zone by monsoon-driven upwelling, and therefore confined close to upwelling centres during the SWM (Krey & Babenerd, 1976; Qasim, 1977, 1982). The JGOFS results have shown that regional and seasonal changes in PP are much smaller than previously believed (e.g. Fig. 6.4; Barber et al., 2001; Smith, 2001; Marra, 2002). The average PP along the US JGOFS southern transect varied within a narrow range of 1.03–1.64 g C m−2 d−1 with identically high values recorded during the peaks of NEM and SWM (Table 6.1). Broadly consistent with these results, the NIOP (Netherlands Indian Ocean Program) cruises in 1992–93 (Veldhuis et al., 1997) also yielded rather close estimates for average PP for the SWM (1.25 g C m−2 d−1) and the NEM (0.80 g C m−2 d−1). However, other JGOFS investigators (e.g. Savidge & Gilpin, 1999) reported a broader range of values and seasonal means. Note that the US JGOFS survey did not cover the most oligotrophic pre-SWM period [April/May; the SI cruise was conducted when the remnants of winter cooling were still discernible (Figs 6.3b and 6.3e)] when high SST and low phytoplankton biomass inferred from satellite imageries are indicative of lower PP (Figs 6.2b and 6.2e). Finally,

200

20 S03 S07 15

S11 S15

10

150

55

60 65 Longitude (E)

70

100 50 SO2 0 0

200

SO4

SO7

400 600 Distance offshore (km)

S11 800

1000

Fig. 6.4 Primary productivity, integrated to the 1% light depth, along the US JGOFS southern transect for NEM (circles), SI (squares) and SWM (triangles). Station locations are shown in the inset. Redrawn from Barber et al., 2001 with permission from Elsevier Science.

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Table 6.1 Cruise-averaged primary production and surface chlorophyll along the southern transect of the US JGOFS Arabian Sea Process Study (after Barber et al., 2001) Northeast Spring Mid Southwest Late Southwest Northeast Monsoon Intermonsoon Monsoon (July/ Monsoon Monsoon (January 95) (March/April 95) August 95) (August/ (December 95) September 95) Primary productivity 1.64 ± 0.16 (g C m−2 d−1) 0.48 ± 0.08 Surface Chlorophyll a (mg m−3)

1.03 ± 0.06

1.62 ± 0.12

1.32 ± 0.13

1.06 ± 0.05

0.15 ± 0.02

0.58 ± 0.07

0.63 ± 0.08

0.45 ± 0.03

the annual cycle of PP deduced from chl a, surface irradiance and MLD data recorded by the Office of Naval Research (ONR) mooring (Dickey et al., 1998) deployed near US JGOFS Sta. S7 during 1994–95 also exhibits considerable seasonality with the maxima and minima occurring during the two monsoons and intermonsoons, respectively (Marra et al., 1998). Superimposed over the general bimodal distribution of PP one can also see variations resulting from mesoscale features: both chl a and PP were greatly enhanced by an eddy and an upwelling filament that passed over the ONR mooring in December 1994 and August 1995, respectively (Marra, 2002). The brief periods of low PP notwithstanding, the annual mean productivity of the Arabian Sea is similar to that of the North Atlantic during spring blooms (Barber et al., 2001). Going on the satellite data, seasonal changes in PP appear to be more pronounced in the eastern Arabian Sea in comparison to the western region (Fig. 6.2). Note, however, that remotely sensed chl concentrations are likely to be overestimated for the coastal zone off India due to the presence of organics (yellow substances) and particulate matter transported by land runoff which obscure the ocean colour signal. Although the cold upwelled water, having >20 μM NO3− to begin with, is generally confined to a few metres below the surface by the low-salinity lens, it receives enough sunlight to enable plankton growth (Fig. 6.5). Moreover, as compared to the western Arabian Sea, upwelling velocities are probably much smaller in this region ensuring a more effective utilization of nutrients. Consequently, chl a and PP levels may reach up to 10 mg m−3 and 6 g m−2 d−1 respectively off the Indian coast. The modest throughput of water also entails greater O2 depletion culminating in anoxic conditions, as we shall see later. The loss of NO3− through denitrification is expected to eventually lower PP, but once the system turns anoxic, nitrogen regenerated from organic matter accumulates as ammonia, and its upward diffusion could still support moderate PP (1–2 gm−2 d−1) in the thin oxygenated surface layer. In waters with O2 < 0.1 ml l−1, conditions are probably too harsh for the normal physiological functions of phytoplankton causing a rapid decrease in PP (Fig. 6.5).

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Oxygen (μM)

Temperature (°C) 20 22 24 26 28

0

100

200

Chlorophyll a (mg m–3) 0

4

2

Depth (m)

0

20

40 (a)

(b)

(c)

60 34

35 Salinity

36 0

8

16

Nitrate (μM)

24 0

100

200

PP (mg C m–3 d–1)

Fig. 6.5 Vertical profiles of (a) temperature (- -) and salinity (–); (b) oxygen (− + −) and nitrate (- • -); and (c) chlorophyll a (- • -) and PP (− + −) at a station located over the Indian continental shelf off Mangalore (12°54′ N, 74°11′ E) sampled on 20 September 2001 (Wajih Naqvi, Shailaja & Prabhu Matondkar, unpublished).

The absence of winter convection in the northward flowing warm, fresher WICC during the NEM, which induces downwelling along the Indian coast, leads to the prevalence of oligotrophic conditions and low PP (<0.5 g C m−2 d−1) in the eastern Arabian Sea (Shailaja, unpublished data). The combination of low concentrations of dissolved inorganic nitrogen (DIN) and calm weather creates an ecological niche for nitrogen fixers during the late NEM-early SI. This period is therefore characterized by widespread blooms of Trichodesmium, which probably play an important but hitherto poorly quantified role in regional biogeochemistry (Qasim, 1970; Devassy et al., 1978). Measurements of N-fixation rates and natural abundance of N-isotopes in NO3− have led to the suggestion that N-fixation in this region may be more important than elsewhere in the ocean (Dugdale et al., 1964; Brandes et al., 1998; Capone et al., 1998). 6.4.3

New production

The large-scale nutrient supply to surface waters during the monsoons implies high rates of new production (NP) especially in regions and during periods of upwelling and vertical mixing. In accordance with these expectations, Owens et al. (1993) found the f-ratio (NP/PP, determined from uptake of 15N-labelled NO3− and NH4+) to increase from 0.1 at the equator to 0.9 in the Omani upwelling region during a study undertaken in September–October 1986. They concluded that the NP could account for approximately half of the PP. Their data also exhibited the expected positive correlation between the f-ratio and PP. However, subsequent JGOFS observations have revealed a more complex

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Table 6.2 Seasonally averaged f-ratios for various sub-regions of the US JGOFS study area (coastal upwelling: Stas. N1–N2, S1–S4; Somali Jet axis: S5–S10; southeast: S12–S15; off Indian coast: N4–N10) – after Sambrotto (2001)

Spring intermonsoon Southwest monsoon

Coastal upwelling

Somali Jet axis

Southeast

Off Indian coast

0.17 ± 0.11 0.22 ± 0.12

0.11 ± 0.05 0.11 ± 0.04

– 0.16 ± 0.11

0.30 ± 0.06 0.10 ± 0.07

pattern: Along the ARABESQUE transect of the UK JGOFS, worked during November–December 1994 (D212), Watts and Owens (1999) not only observed lower f-ratios (0.2–0.52) with less distinct onshore–offshore gradients, but the depth-integrated ratio was also found to be inversely related to the mixed-layer NO3− inventory with different relationships for the stations located north and south of Latitude 17°30′ N. The latter was interpreted to reflect the existence of different biogeochemical provinces (see also Brock etal., 1998; Watts etal., 1999). Sambrotto’s (2001) analysis of the US JGOFS data for the SI (TN045, March– April 1995) and early SWM (TN049, July–August 1995) cruises yielded even lower (<0.05) f-ratios at stations located closest to the coast along the southern transect, increasing to peak values (>0.2) at the edge of the upwelling zone during both seasons. As a result, the average f-ratio for the upwelling zone was moderately high but not the maximum among all sub-regions (Table 6.2). This has been attributed to the preferential uptake of NH4+ and its inhibitory effect on NO3− uptake (Watts & Owens, 1999; Sambrotto, 2001). Nevertheless, Sambrotto (2001) also observed rapid NO3− uptake in upwelling filaments, and proposed that offshore transport of nutrient rich water with established phytoplankton populations might compensate for the degraded light conditions resulting from deep mixed layers, thereby sustaining high rates of NP. 6.4.4

Phytoplankton composition and size distribution

The highly variable physical forcing and the resultant hydrographic changes in the Arabian Sea are expected to give rise to widely differing phytoplankton assemblages in both space and time. As in other areas, pico-plankton are expected to be the dominant contributors to PP in areas where, and during periods when, oligotrophic conditions prevail in the euphotic zone. Conversely, phytoplankton community structure is expected to shift to larger cells when and where the surface waters are nutrient replete. These trends are confirmed by the results of numerous recent studies following a variety of approaches based on measurements of size-fractionated 14C uptake and/or chl a (Owens et al., 1993; Savidge & Gilpin, 1999), chemotaxonomic phytoplankton pigments (Veldhuis et al., 1997; Latasa & Bidigare, 1998; Barlow et al., 1999) and microscopic and/or flow cytometric analyses (Burkill et al., 1993; Veldhuis et al., 1997; Campbell et al., 1998; Garrison et al., 1998, 2000; Brown et al., 1999; Dennett et al., 1999; Tarran et al., 1999; Shalapyonok et al., 2001).

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Distributions of 14C PP among three size groups of phytoplankton (0–2, 2–18 and >18 μm) during the two ARRABESQUE cruises (D210 and D212) are shown in Fig. 6.6 (Savidge & Gilpin, 1999). Although D210 was undertaken in September – towards the end of the SWM – nutrient enrichment in surface waters extended up to 1000km offshore (Woodward etal., 1999). Substantial small-scale variations in physical and chemical properties, reflecting complex small- and meso-scale dynamics of water masses, characteristic of the Arabian Sea, were noticed. The size-fractionated chl a and PP data also showed this variability superimposed over an overall trend of dominance of production by autotrophs in the >18 μm fraction within the coastal upwelling zone (where the columnintegrated PP and chl a values reached up to 3.8 g C m−2 d−1 and 69 mg m−2, respectively) and by those in the 0.2–2μm fraction offshore (where PP and chl a averaged around 1.5g C m−2 d−1 and 30mgm−2, respectively). The geographical variability was much smaller during the early NEM (when the average PP decreased to around 0.75 g C m−2 d−1, and the chl a inventory was at a modest ASS A1 3000

A2

A3

A4

A5

2500

20° N

2000 10°

1500 1000

A7 (a)

Oman AS5 AS3 A1A2 A3 A4 A5 A6 A7 60° E

70°

500

Production (mg C m

–2

–1

d )

A6

AS2

0 1000

AS3

(b) >18 2–18 0.2–2

800 600 400 200 0 20

17.5

15

12.5

10

7.5

Latitude (N)

Fig. 6.6 Distribution of size-fractionated primary production along the ARABESQUE transect during (a) late SWM (D210) and (b) early NEM (D212). Station locations are shown in the inset. Redrawn from Savidge and Gilpin (1999) with permission from Elsevier Science.

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level of 20 mg m−2) with most of the production occurring through organisms in the pico-size range (0.2–2 μm) all along the transect (Fig. 6.6). The relative uniform geographical distributions of biological variables reflected the small onshore–offshore gradients in surface nutrients, as the mixed layer concentration of NO3− rarely exceeded 1 μM during this season (Woodward et al., 1999). Phytoplankton abundance and biomass, determined by microscopy and flow cytometry, during the same cruises are consistent with the results of sizefractionated PP (Tarran et al., 1999). In general, pico-planktonic prokaryotes, Synechococcus and Prochlorococcus dominated the autotrophic assemblages, and of the nano- and micro-autrotrophs, only diatoms contributed significantly to the phytoplankton biomass inshore of Sta. A3 during the SWM (Fig. 6.7). The most abundant diatom genera there were Corethron, Porosira and Rhizosolenia. Prochlorococcus was restricted only to the offshore region during the SWM, but was abundant everywhere during the early NEM, dominating the phytoplankton SW Monsoon

Early NE Monsoon

AS5-1

AS5-1 Synecho

AS2-3

A1-8

AS4-1

Prochloro picoeuc Nanoflag Cocco Dinoflag Diatom

AS3-4

Upwelling region

A1-30

AS2-1 AS1-1 AS1-9 AS1-23 AS1-36

A1-42

A2-4

AS2-4

A3-23

AS3-6

A3-41

AS3-20

Station

AS1-36

Transitional region Station

AS3-2

A3-26 A3-35

Downwelling region

A3-47 A4-2 A5-8

A4-2 A5-6

A6-2

A6-1

A7-5

A7-6

Oligotrophic region

A7-30

A7-19 A7-126

A7-41 10

8

6

4

2

0

2

4

6

8

10

Biomass (mg C m–2 X 1000) Fig. 6.7 Standing stocks of various phytoplankton, integrated to 80 m, along the ARABESQUE transect during two seasons. Reproduced from Tarran et al. (1999) with permission from Elsevier Science.

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biomass at Stas. A6 and A7; apparently its abundance was favoured by high temperature (≥27°C) and low NO3− (<0.1 μM) (Tarran et al., 1999). In contrast, Synechococcus was quite abundant, very often predominant, during both the seasons except in the most oligotrophic offshore waters (Stas. A6 and A7) during the early NEM. The above trends are generally supported by the results of the corresponding US JGOFS cruises (TN050 – August–September 1999 and TN054 – December 1995). While some inter-annual changes could be recognized [e.g. the phytoplankton stocks were lower on TN050 (1–3 g m−2, Garrison et al., 2000) than on D210 (1–9 g m−2, Tarran et al., 1999)], these could arise in part from sampling bias and/or different conversion factors used to compute biomasses from cell volumes or counts. The latter probably explains why picoeukaryotes, which were regarded as relatively insignificant contributors to phytoplankton biomass by Tarran et al. (1999), were credited for a large portion of PP in the coastal upwelling zone by Campbell et al. (1998) and Brown et al. (1999). More definitely, as compared to D210, high abundances of nano- and microautotrophs extended much farther offshore on TN050 when blooms of the prymnesiophyte Phaeocystis were found at Stas. S7 and S11 (Garrison et al., 1998). During the late NEM (January–February) of 1995, sampled on TN043 of US JGOFS, nutrient enrichment associated with convective mixing was found to result in an increase in the average size of the autotrophs (due to greater abundances of diatoms and autotrophic nanoplankton) but to a smaller degree than the SWM. While, as will be discussed later, low euphotic zone silicate (SiO44−) concentration probably limits diatom growth during this season, colder SSTs and moderate NO3− concentrations are not conducive for the growth of Prochlorococcus, explaining its very low abundance except in the most offshore waters (Campbell et al., 1998). The Synechococcus biomass was found to be moderate and relatively invariant geographically. In contrast, oligotrophic conditions during the SI (TN045) were associated with the predominance of Prochlorococcus all over the Arabian Sea with the maximum counts (7 × 105 cells ml−1) being the highest recorded anywhere in the ocean. Synechococcus was less abundant as compared to the two monsoon seasons (Campbell et al., 1998). The biomass of microplankton (mostly diatoms) was expectedly low except at the stations along the northern transect which were probably sampled when remnants of convectively injected NO3− were still present in the surface layer (Morrison et al., 1998). In the upwelling zone off Somalia, surveyed by the NIOP during the SWM (May–August 1992) as well as the NEM (January–February 1993), diatoms (mostly belonging to genera Nitzschia, Rhizosolenia and Coscinodiscus) were the predominant autotrophs; Synechococcus and picoeukaryotes were also present in significant numbers, but Prochlorococcus was absent (Reckermann & Veldhuis, 1997; Veldhuis et al., 1997). Phytoplankton composition during the

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NEM, on the other hand, was characterized by the dominance by pico-plankton, particularly Synechococcus, with lesser contributions from Prochlorococcus and picoeukaryotes (Veldhuis et al., 1997). 6.4.5

Chlorophyll and POC

Concentrations of chl a and POC exhibit considerable variability in both space and time as exemplified by vertical sections of these properties, based on observations made with CTD-mounted sensors of fluorescence and light transmission along the US JGOFS southern transect (Fig. 6.8). Even though the fluorescence and beam attenuation data were calibrated with discrete measurements of chl a and POC, the profiler-derived values differed randomly, and sometimes systematically, from actual concentrations due to substantial scatters in the relations used and the assumptions involved (Gundersen et al., 1998). As an example, for TN050 the in situ florescence-based chl concentrations are significantly higher than those measured for the discrete samples (Table 6.1). However, both sets of data show remarkably similar

Fig. 6.8 Distributions of chlorophyll fluorescence (panels on the left) and POC (panels on the right) in the upper 150 m along the US JGOFS southern transect during different seasons. These properties were derived from in situ light transmission and fluorescence measurements. The dotted and solid lines indicate the mixed layer depth (identified by an increase in σθ by 0.03 relative to the surface) and the 0.5 μM NO3− contour, respectively. Reproduced from Gundersen et al. (1998) with permission from Elsevier Science (see Color Plate 3).

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

Stations S3

S6

S7

S11

S13

S14 S15

1250

1450

0.50

1.00

60

S9

1.00

1.00

40

S8

0.50 1.00 1.50

50 0.

20 Depth (m)

S4 S5

1.00

0

S1 S2

0.50 0.50

80 100 50

250

450

650

850

1050

Distance offshore (km) Fig. 6.9 Distribution of chlorophyll a (mg m−3), measured in discrete samples, in the upper 100 m along the US JGOFS southern transect during the late SWM (TN050). Reproduced from Gundersen et al. (1998) with permission from Elsevier Science.

trends (Figs 6.8 and 6.9). The most striking aspect of the chl distribution is the high concentrations in the surface layer during both the SWM and NEM seasons. In general, the highest concentrations occur in the Omani coastal waters, although patches of very high chl are also encountered offshore, presumably associated with the mesoscale features such as filaments and eddies. In conformity with the satellite data, elevated chl levels persist ≥1000 km from the Omani coast. During the SI, chl levels are quite modest in the shallow surface mixed layer, but since the thermocline extends into the euphotic zone a subsurface chl maximum (SCM) develops just below the mixed layer. However, since the irradiance due to chl received by the satellite is integrated down to a depth of 1/k, and k, the vertical attenuation coefficient of light, ranges between 0.03 and 0.1 m−1 in the Arabian Sea, the SCM is beyond the reach (10–30 m) of the satellite-borne sensors (Banse, 1987; Gundersen et al., 1998). The presence of the SCM dampens the magnitude of seasonal variability of depth-integrated chl a inventory. The same is also true for POC, the trends of variability of which are quite similar to those of chl (Fig. 6.8). Note, however, that at the SCM the PP/chl a ratio is expected to be considerably lower than at the surface due to the effect of shading (Banse, 1994), and so even in cases where the SCM is accompanied by a subsurface production maximum, the columnintegrated PP is only moderately high (Pollehne et al., 1993). 6.4.6

Effect of changes in mixed layer depth

The PP rate and related chl a and POC concentrations are controlled to a very large extent by the mixed layer dynamics, as can be seen in Fig. 6.8, but the

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relationship between the MLD and PP is complex, and is variable seasonally as well as regionally. At the ONR mooring site, for instance, the shallowest mixed layer was found to be associated with the lowest chl a and PP during the SI (Marra et al., 1998). The chl a concentrations rose in concert with windinduced mixed layer deepening during the early part of the SWM (June–July), but the highest chl a and PP were recorded towards the end of the season (September) after the re-establishment of water column stratification (Marra et al., 1998). The deepest mixed layers in the region occurred during the NEM due to convective overturning (Banse & McClain, 1986; Banse, 1987) and were associated with elevated chl a (Fig. 6.8) and PP in spite of the expected shorter exposure time of phytoplankton to light. It has been suggested that solar heating during the day may bring about significant diurnal oscillations in MLD, affecting both the phytoplankton growth and nutrient recycling (Wiggert et al., 2000; Wiggert, 2002). In the eastern Arabian Sea, PP is inversely correlated with the MLD (Matondkar, unpublished data). Very shallow mixed layers (<10 m) in summer and autumn probably allow the satellite to partially sense high subsurface chl over the inner shelf in this region. Using a biophysical model consisting of four active layers (surface mixed layer, diurnal thermocline layer, seasonal thermocline and main thermocline) that was forced by monthly climatological and daily fields, McCreary et al. (2001) found the surface mixed layer to deepen steadily (entraining nutrients from the thermocline) through the two monsoon seasons in solutions forced by climatological fields. In this case, intense, short-lived phytoplankton blooms developed towards the end of both SWM and NEM (spring and fall blooms) due to detrainment and the consequent increase in light available to phytoplankton in the mixed layer. In model runs utilizing daily fields, detrainment occurred during monsoon breaks causing more frequent blooms. Diurnal forcing (incorporation of diurnal irradiance cycle) affected the biophysical responses by lengthening the duration of the deep mixed layer during the NEM, by strengthening the spring and fall blooms and delaying them by three weeks, and by intensifying phytoplankton levels during the intermonsoon periods. The best correspondence to the phytoplankton production time series of Marra et al. (1998) was obtained with the solution driven by daily fields with diurnal forcing (McCreary et al., 2001). However, as pointed out by Sathyendranath and Platt (1994) and Sathyendranath et al. (1999), models that only consider mixed layer dynamics without taking into account changes in bio-optical properties of phytoplankton are probably insufficient to explain the occurrence of phytoplankton blooms in the Arabian Sea. Sathyendranath et al. (1999) provided evidence for significant seasonal changes in the parameters that describe the dependence of photosynthesis rate of phytoplankton on the amount of light available (P–I curves), viz. the assimilation number (PmB) and the initial slope of photosynthesis-light curve (αB), as well as in the specific absorption coefficient of phytoplankton (a c*).

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

Contrary to anticipation and for reasons not entirely clear, both PmB and αB, expressed in units of mg C (mg chl a)−1 and mg C (mg chl a)−1 h−1 (μE m−2 s−1 )−1, respectively, were on average lower during the SWM (D210) (4.15 and 0.023) relative to the early NEM (D212) (5.43 and 0.028) and the late NEM (NIOP) (5.42 and 0.033). But the effect of these decreases in P–I parameters on PP was more than compensated by lower ac* [0.020 during the SWM against 0.038 and 0.031 m−2 (mg chl a)−1 during the early and late NEM, respectively], presumably arising from seasonal changes in phytoplankton composition (i.e. bigger cells – mostly diatoms – during the SWM absorbing light less efficiently). Thus, the maximum quantum yield for photosynthesis may be enhanced substantially during the SWM, contributing to the initiation of phytoplankton blooms. Of course, in addition to the above-mentioned factors that regulate the growth of phytoplankton, the occurrence of blooms is also critically dependent on the processes that control phytoplankton loss from surface waters. We shall describe these processes in some detail in the following sections. 6.5 6.5.1

Heterotrophic biomass and production Heterotrophic bacteria

There have been several studies of bacterial abundance and productivity in the Arabian Sea (Ducklow, 1993; Ramaiah et al., 1996; Wiebinga et al., 1997; Campbell et al., 1998; Garrison et al., 2000; Pomroy & Joint, 1999; Ducklow et al., 2001). These have revealed that the heterotrophic bacteria, expectedly most abundant in the surface layer, numerically dominate microbial population; also, their abundance in the region (>5 ×108 cells l−1) is generally higher throughout the year as compared to other tropical regions. The US JGOFS surveys have enabled an assessment of the seasonal and geographical changes in bacterial abundance and productivity in the western and central Arabian Sea (Campbell et al., 1998; Garrison et al., 2000; Ducklow et al., 2001). Except for the early NEM (TN054), near-coast enhancement of bacterial abundance/biomass (measured by epifluorescence microscopy and flow cytometry), was not noticed (Fig. 6.10). Elevated surface counts (>15 × 108 cells l−1) and vertically integrated biomass did occur during the late SWM (TN050), but these were spread over a much larger area, extending beyond the zone affected by upwelling (Ducklow et al., 2001). Surprisingly, nutrient enrichment through vertical mixing did not enhance the bacterial numbers during the late NEM (TN043); in fact, the lowest and geographically invariant surface values (5–10 × 108 cells l−1) were recorded during this season. In contrast, the SI was characterized by some of the highest bacterial counts at the surface (reaching up to 20 × 108 cells l−1); the subsurface abundance maximum was also more pronounced during this season, resulting in high vertically integrated biomass. The euphotic zone bacterial production was estimated to be 187, 262,

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10

10 15

10

5

December

5

10

15

no data

20

10

10

10

20 10 0

15

no data

–1

10

10

15

V

1520

15 20

10

5

10 60

65

70

60

65

70

60 65 70 East longitude

60

65

70

60

65

Flow cytometry 108 cells l–1

North latitude

10

August

July

Microscopy 108 cells l–1

20

March

10

January

70

Fig. 6.10 Distributions of bacterial abundance at the sea surface during various seasons. Two different methods – epifluorescence microscopy (upper panels) and flow cytometry (lower panels) – were used to enumerate bacteria. Redrawn from Ducklow et al. (2001) with permission from Elsevier Science.

353 and 146mgC m−2 d−1 for TN043, TN045, TN049 and TN054, corresponding to 18, 26, 24 and 15% of PP. Thus, although the average bacterial production appears to be at a seasonal maximum during the SWM, consistent with the observations off Somalia (Wiebinga et al., 1997), there is no evidence for bacterial blooms during this season (Ducklow et al., 2001). The low amplitudes of oscillations in bacterial biomass and production, probably reflect the abovementioned small seasonality and onshore–offshore gradients in PP (Barber et al., 2001), and in dissolved organic carbon (DOC) concentration in the surface layer (Hansell & Peltzer, 1998). 6.5.2

Nano- and microheterotrophs

Heterotrophs in the nano (2–20 μm) and micro (20–200 μm) size ranges, largely comprising protozoa and micrometazoa, are an important component of the pelagic food web (Azam et al., 1983). Their importance in carbon cycling owes to their ability to consume small particles such as the pico- and nano-plankton in the oceanic surface layer, and also to their high respiratory activity which favours mineralization of organic matter in the surface layer over its export to the deep sea (Garrison et al., 2000). Their distribution in the Arabian Sea has been recently investigated by several JGOFS teams (Gauns et al., 1996; Reckermann & Veldhuis, 1997; Garrison et al., 1998; Dennett et al., 1999; Stelfox et al., 1999). Like bacteria, heterotrophic nano- and microplankton have also been found to be predominantly concentrated in surface waters. The observed regional and temporal changes for both groups are not very pronounced, barring generally higher microzooplankton values reported by Gauns et al. (1996) from the eastern

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and central Arabian Sea, and in general, their biomasses are of similar magnitude (Fig. 6.11). For instance, the integrated nano- plus microhetreotrophic biomasses in the upper 100m in the US JGOFS study area are 1.36, 1.48, 1.53 and 1.27g Cm−2 for TN043, TN045, TN050 and TN054 respectively (Garrison et al., 2000), accounting for 36–57% of the total microbial biomass. As mentioned earlier,

30

(a)

Biomass (mg C m–3)

S.W. Monsoon

20

10

S2 A1 S4 A2 A3 S7

S15 A6

30

(b)

Biomass (mg C m–3)

Early N.E. Monsoon

20

10

0 S2 A1 S4 A2 A3 S7

S15 A6

Fig. 6.11 Biomasses within different size fractions at comparable stations sampled by the US JGOFS (S2–S15) and ARABESQUE (A1–A6) investigators during (a) SWM and (b) NEM (see insets of Figs 6.4 and 6.6 for station locations). For each station there is a set of three bars showing contributions by pico (filled), nano (hatched) and micro (open) plankton. The bar in the middle is for autotrophs and the other two for heterotrophs. Reproduced from Garrison et al. (2000) with permission from Elsevier Science.

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nutrient enhancement increases the size of autotrophs; these are usually accompanied by increases in the nano- and microheterotroph sizes, indicating a dynamic food web (Garrison et al., 2000). 6.5.3

Mesozooplankton

Mesozooplankton are the major consumers of organic matter synthesized by phytoplankton and re-packagers of small particles. As a result, they form a critical component of the oceanic “biological pump”. General trends of their variability in the Arabian Sea have long been known, especially after the International Indian Ocean Expedition (IIOE) (e.g. Vinogradov & Voronina, 1961; Rao, 1973; Smith, 1982; Madhupratap & Haridas, 1986; Madhupratap et al., 1990; Banse, 1994; Bottger-Schnack, 1996). Systematic observations made under JGOFS have provided important insights into their role in biogeochemical cycling (Madhupratap et al., 1996a,b, 2001; van Couwelaar, 1997; Smith et al., 1998; Wishner et al., 1998; Stelfox et al., 1999; Smith, 2001, 2002). One outstanding feature of the distribution of mesozooplankton biomass in the Arabian Sea, known as the Arabian Sea paradox, is its general geographical and seasonal invariability (e.g. Baars & Brummer, 1995; Madhupratap et al., 1996a; Baars & Oosterhuis, 1998). For instance, peak zooplankton biomass in the upper 200 m reported from the relatively oligotrophic eastern Arabian Sea (3.4 g m−2, Mathew et al., 1990) is 79% of the highest value reported from the upwelling zone off Somalia (4.4 g m−2, Baars & Oosterhuis, 1997) and 65% of the maximum off Oman as well as for the entire Arabian Sea (5.33 g m−2, Qasim, 1977). It may be argued that since the earlier estimates were based on collections using nets of a larger mesh size (330 μm) than that used during the JGOFS (200 μm), except on the NIOP cruises, any variability in the 200–330 μm could have been missed. But elevated biomasses have also been recorded with the finer net in distinctly oligotrophic areas implying a different food web (e.g. the microbial loop, Madhupratap et al., 1996a, 2001). Alternatively, as pointed out by Smith (2001), the paradox may simply reflect relatively uniform food supply arising from small changes in PP discussed above. Nevertheless, as demonstrated by Wishner et al. (1998), some fluctuations in mesozooplankton biomass do occur from coastal to offshore waters and from one season to another (Fig. 6.12). The biomass was at maximum in waters closest to the coast both in the upper 350 m (Smith et al., 1998) and 1000 m (Wishner et al., 1998), and decreased generally in the offshore direction. The onshore–offshore gradient was the strongest during the late SWM and the weakest during the NEM. The period of maximum biomass was not the same as that of the highest PP (SWM) at all stations – the biomass peaked during the SI in offshore waters. The onshore–offshore trends in biomass in various size fractions (202–500, 500–1000, 1000–2000 and >2000 μm) were complex, but the largest fraction generally accounted for most of the biomass. In the upper 200 m, where

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Fig. 6.12 Distribution of the total night-time biomass of mesozooplankton in the upper kilometre along the US JGOFS southern transect during the late NEM (crosses), SI (triangles), late SWM (filled circles) and early NEM (open circles). Data from Sta. N7 (19° N, 67° E) along the northern transect are also included. Reproduced from Wishner et al. (1998) with permission from Elsevier Science.

most of the biomass seems to reside, biomass in the smallest size fraction was an order of magnitude greater during the late SWM and the early NEM than during the late NEM. For the largest size fraction in the same depth range, on the other hand, the NEM was the season of the highest biomass especially in offshore waters. The seasonal variability was more pronounced for the largest size fraction in offshore waters relative to nearshore waters (Wishner et al., 1998). Mesozooplakton biomass in the upper 300 m recorded by Stelfox et al. (1999) on the two ARABESQUE cruises was relatively lower and the onshore– offshore gradients steeper than those for the corresponding US JGOFS cruises: on D210, the average values for the upwelling, mesotrophic downwelling and oligotrophic zones were 1.25, 0.61 and 0.21 g m−2 respectively, as compared to 1.98 g m−2 at Sta. S2 and 1.03 g m−2 at Sta. S15 on TN050. The values decreased by about half on D212. For the same depth range in the Somali upwelling zone, the NIOP observations yielded an average biomass of 1.32 g m−2 during the SWM and 1.07 g m−2 during the NEM (van Couwelaar et al., 1997), not very different from other values from this region reported earlier (Cushing, 1973). In comparison, Madhupratap et al. (2001) reported a broad range of values (0.86–3.3 g m−2) for the 0–500 m depth range along the JGOFS-India transect (3–21° N latitude, 64° E longitude) during September–October 1992, 1993. Vertical distributions of several zooplankton parameters such as the biomass, taxonomic composition and size exhibit considerable seasonal and regional variations. Rapid decreases in the biomass (by one to two orders of magnitude) occur within the thermocline. When the subsurface O2 concentration falls below 0.1 ml l−1, the extent of vertical migration, and hence the difference between

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the day- and night-time biomasses is greatly suppressed; however, some zooplankton such as ostracods (Madhupratap et al., 2001) and the copepod Pleuromamma indica (Saraswathy & Iyer, 1986) appear to be capable of tolerating suboxic conditions and therefore migrate vertically deep into and out of the suboxic zone (Smith et al., 1998; Wishner et al., 1998; Madhupratap et al., 2001; Smith, 2001). In the area investigated by US JGOFS, the greatest day/night contrast in biomass was observed in the nearshore waters. However, the pattern of migration changed seasonally: the normal migration pattern (night-time biomass in the surface layer being higher than the day-time biomass) reversed during the SWM, most likely because of the presence of copepods Calanoides carinatus and Eucalanus subtenuis, which dominate the biomass in the upwelling zone and do not exhibit diurnal vertical migration (Smith et al., 1998). The zooplankton composition differs substantially from place to place and from season to season. For example, while the small calanoid copepods have been found to be predominant during the SWM, cyclopoid copepods are most abundant during the NEM (Madhupratap et al., 1996a; Smith et al., 1998). Groups other than copepods also make significant, sometimes dominant, contribution to the zooplankton biomass. For example, along the JGOFS-India transect during the FI, Madhupratap et al. (2001) observed the dominance of copepods up to 15° N latitude, north of which ostracods became equally or even more important. Overall, the two groups accounted for an average of 85% of the total zooplankton biomass with the remainder made up mostly by chaetognaths, euphausids, siphonophores and tunicates. 6.6

Food web structure and export of material to the deep sea

As described above, large hydrographical changes greatly influence the phytoplankton community structure and hence the functioning of the biological pump in the Arabian Sea. As in other parts of the oceans, oligotrophic waters with pico-plankton as the dominant primary producers are expected to be characterized by phytoplankton growth rates that are substantially less than the physiologically maximal values; grazing by microzooplankton is expected to more or less balance the growth, efficiently remineralizing nutrients, such that most of the PP is regenerated with little export of carbon to the deep sea. In contrast, addition of new nutrients is expected to support the growth of large phytoplankton at rates close to the physiological maxima and greater utilzation by mesozooplankton. A large fraction of PP in such environments is new and exported out of the surface layer. The following results generally conform to these anticipations. 6.6.1

Phytoplankton growth and mortality

Measurements of rates of phytoplankton growth and mortality through microzooplankton grazing made during the US JGOFS expedition [on TN050 and

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TN054 by Landry et al. (1998) and on TN043 and TN045 by Caron and Dennett (1999)] revealed higher growth rates during the productive seasons (averaging 0.79 d−1 for TN050 and 0.84 d−1 for TN043) than those for the less productive ones (averaging 0.66 d−1 for TN054 and 0.44 d−1 for TN045). The mortality rates averaged 0.55, 0.57, 0.35 and 0.30 d−1 for TN050, TN054, TN043 and TN045 respectively. In both studies, the growth rates were higher at stations where surface NO3− concentrations exceeded 1 μM than those with lesser surface NO3− (e.g. 1.1 d−1 versus 0.5 d−1 for TN050 and TN054), such that mortality nearly balanced growth at the oligotrophic sites (generally on TN045) while a net growth occurred in waters having sufficient NO3− (generally on TN043). However, in conflict with these observations, Edwards et al. (1999) found the growth rate to increase from 0.2 d−1 nearshore to 1.8 d−1 offshore during the SWM (D210) in spite of the inability of microzooplankton to effectively graze on diatoms. The growth rates during early NEM (D212) were fairly uniform across the basin, averaging 0.68 d−1. The cruise averages for the mortality rates were 0.33 d−1 for D210 and 0.41 d−1 for D212. Interestingly, despite the observed differences in geographical trends, the cruise-averaged growth rates varied within a narrow range with the exception of TN045. It can be concluded from these results that, in general, mortality due to microzooplankton grazing is about half of the phytoplankton growth with the difference seemingly reflecting losses through mesozooplankton grazing and sinking. The former, quantified by Roman et al. (2000) for the same US JGOFS cruises, is probably more significant in the Arabian Sea than in other areas due to high mesozooplankton biomass, abundance of diatoms and warm temperatures. According to these authors, on average, mesozooplankton consumed 0.53 g C m−2 d−1, corresponding to ~40% of PP, most of it for meeting their high metabolic requirement. The average growth rate was computed to be 0.14 d−1, and since the mesozooplankton biomass did not vary much with time, it was concluded that the production (0.16 g C m−2 d−1) was more or less balanced by losses through grazing by organisms belonging to the higher trophic level (Roman et al., 2000).

6.6.2

Particle fluxes to deep sea

Sinking flux of particulate matter and its composition have been monitored since 1986 through deployment of sediment traps at three sites, one each in the western, central and eastern Arabian Sea, under an Indo-German collaborative project (Nair et al., 1989; Haake et al., 1993; Rixen et al., 1996, 2002). The results of this study have been supplemented by measurements during 1992–97 under national JGOFS programmes of the Netherlands (Van Veering et al., 1997; Koning et al., 1997; Broerse et al., 2000), Germany (Suthhof et al., 1999), India (Sarin et al., 1996) and US (Lee et al., 1998; Honjo et al., 1999).

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The most striking feature of all moored sediment trap records is the bimodal seasonal distribution with maxima in total flux as well as its various components occurring during the two monsoons; exports to the deep sea are thus tightly coupled to surface oceanographic and meteorological processes (Fig. 6.13). Export fluxes are higher during the SWM in comparison to the NEM especially in areas affected by upwelling (Fig. 6.13), with the SWM flux also determining the magnitude of the inter-annual variability (Haake et al., 1993; Rixen et al., 1996). However, peak fluxes during the two seasons are comparable in the northern Arabian Sea due to the greater influence of winter convection in euphotic zone nutrient enhancement (Suthhof et al., 1999). Mesoscale features such as eddies and filaments have been found to bring about spectacular enhancement of particle sedimentation (Honjo et al., 1999). As in other oceanic areas, POC fluxes in the Arabian Sea also decrease rapidly with depth especially in the upper water column, and decline more slowly with distance offshore (Lee et al., 1998). Much of the organic production is recycled within the surface layer: the US JGOFS data yielded an average of 6.3% of overhead PP passing through the 100 m level via sedimentation (decreasing from 8.2% at the most inshore trap site to 4.3% at the most offshore one). Variations in the accumulation rate of organic matter in sediments were found to be much larger (lower by a factor of 40 in the offshore than in the nearshore regions). Given the relative constancy of both PP and zooplankton biomass in space and time, the range of geographical and temporal changes in POC export is surprisingly large. The other intriguing observation concerns the timing of accelerated exports: marked increases in particle sedimentation occur quickly (within 1 month) after the onset of the SWM, but the chl concentrations in the surface layer do not rise greatly until August (Banse & English, 2000). Nevertheless, maximal fluxes still occur towards the end of the SWM (Fig. 6.13; also see Rixen et al., 2002). Of the various components of the biogenic sinking material, carbonate, comprising tests of coccolithophores and foraminifera, is the first to reach sediment traps; the peak in biogenic opal (frustules of diatoms dominated by Rhizosolenia) occurs in late August/early September (i.e. at least a month after the carbonate peak) (Haake et al., 1993). The increased opal flux raises the Si/Ca ratio in sinking matter (Fasham et al., 2001). The delayed opal sedimentation has been attributed to the low ratio of silicate (SiO44−) to NO3− concentrations in the upwelled/surface-entrained deep water early during the SWM, which can support the growth of organisms with calcareous skeletons but not those with siliceous shells, and a shift to diatom dominated community when the SiO44−/NO3− ratio rises sufficiently (Morrison et al., 1998) to support the growth of diatoms (Haake et al., 1993). Although high flux events are almost always coincident with the dominance of diatoms in surface waters (Buesseler et al., 1998; Garrison et al., 1998; Honjo et al., 1999; Smith, 2001; Rixen et al., 2002), the exact causal link between the diatom blooms and enhanced fluxes is

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Fig. 6.13 Effect of seasonal changes in physical characteristics of surface water, as measured by instruments on a surface air–sea interaction buoy, on particle flux to the deep ocean, recorded by sediment traps at 2 229 m during 1994–95. These moorings were deployed close to the US JGOFS Sta. S7 (Fig. 6.4, inset). Redrawn from Honjo et al. (1999) with the data for the mixed layer depth (MLD, taken as the depth at which temperature decreased by 1°C) from Dickey et al. (1998).

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not fully understood. The diatom population in the offshore mesotrophic waters has been found to be in a senescent state during the late SWM (Garrison et al., 1998), implying that these blooms might have originated near the coast and transported subsequently offshore. The senescent cells might sink either individually or in the aggregated form facilitated, for example, by the mucus produced by Phaeocystis (Garrison et al., 1998). Alternatively, they may be grazed upon by large micro- and mesozooplankton producing dense fecal pallets. However, the mesozooplankton fecal pallet production has been computed to be insufficient to sustain the observed flux during the SWM (Roman et al., 2000). Moreover, under certain circumstances, grazing by mesozooplankton may, in fact, suppress the vertical flux as proposed by Smith (2001) who attributed the large pulse in the flux recorded in early September to the cessation of feeding by Calanoides carinatus and Eucalanus subtenuis. These ontogenetically migrating copepods dominate zooplankton biomass during the period of upwelling when they build up large lipid reserves that enable them to diapause at depth during the non-upwelling period (Smith et al., 1998; Smith, 2001). Smith (2001) argued that their departure from the surface layer toward the end of the SWM would release the predation pressure on diatoms, which would then rapidly consume the remaining nutrients, become senescent and sink out of the surface layer. The pattern of vertical exports from the euphotic zone during the NEM, viz. low fluxes during the period of most intense convection (January–February) followed by a peak close to the transition from winter to spring (March), is counter-intuitive (Fig. 6.13; also see Buesseler et al., 1998; Rixen etal., 2002). The simplest explanation of this anomaly could be that the mixed layer deepening reduces mean underwater irradiance, keeping PP below its full potential until the thermocline shoals up toward the end of the season (Fig. 6.13). This scenario is similar to that of the formation of spring blooms in temperate regions. However, as already stated, daytime heating can stabilize the upper few metres of the water column producing a shallow mixed layer (within the thicker surface layer above the seasonal thermocline) that can allow moderately high PP (Gardner et al., 1999; Wiggert et al., 2000; McCreary et al., 2001). This diurnal cycle, which can account for the incomplete utilization of NO3−, would also facilitate the production of ammonium (NH4+) through mineralization of organic matter in the lower part of the surface layer. The NH4+, when supplied to the euphotic layer as a result of night-time mixing, can inhibit the uptake of NO3− by phytoplankton (McCarthy et al., 1999). Thus, a significant fraction of elevated PP could be regenerated rather than new explaining the modest f-ratios (averaging 0.15) observed by McCarthy et al. (1999) on TN043. It is obvious that while the total PP is quantitatively similar during the NEM and SWM, the quality of production and the food web structure should be very different during the two seasons. Naqvi et al. (2002) argued that a deficiency of SiO44− relative to NO3− (reflecting the concentrations of these nutrients at the

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horizon – 100–120 m – to which the convection extends) may limit the growth of diatoms during the NEM, as was proposed by Haake et al. (1993) for the early SWM period. Diatoms did dominate phytoplankton composition much farther offshore on TN050 (Garrison et al., 1998) than on TN043 (Dennett et al., 1999), and the percent chl in >2 μm fraction averaged 63 and 30 respectively on these cruises (Roman et al., 2000). It seems very likely that the high NO3− and low SiO44− levels favour the growth of smaller autotrophs (e.g. Synechococcus and small eukaryotic algae). This may create an ecological niche for filter feeders such as salps that can efficiently utilize food particles in a fewmicrometre size range. Swarms of salps have been sighted on several occasions in the northern Arabian Sea during the NEM (Naqvi et al., 2002). Through their characteristic indiscriminate grazing, these large planktonic microphages have the potential to remove chl from water thereby moderating PP. However, although salps are also known to produce large fecal pellets that have high sinking rates (Madin, 1982), drifting traps deployed at 140 and 300 m during a salp swarm collected modest quantities of POC (<32 mg C m−2 d−1), implying recycling of the pellets (Naqvi et al., 2002). This is consistent with the results of Roman et al. (2000) who found a similar mismatch between the vertical POC fluxes and estimated production of fecal pallets by zooplankton. Over the period of their time-series study, Naqvi et al. (2002) noticed substantial buildup of DOC in the upper 350 m. How, and in fact whether, this was related to the activity of salps could not be established. However, more comprehensive measurements of DOC made during the US JGOFS have revealed that the DOC does accumulate seasonally (between NEM and SI) in the upper water column in this region (Hansell & Peltzer, 1998). The estimated increase in DOC inventory (31–41 × 1012 g C north of 15° N latitude, amounting to 6–8% of PP) is very similar to the amount of POC vertically exported from the surface layer. According to Lee et al. (1998) about 6% of PP settles through the 100 m horizon as POC, and its absolute amount has been computed to be 84–91 × 1012 g C y−1 for the region bounded by 6° N latitude and 2 km-depth contour (Rixen et al., 2002). Thus, it may be concluded that, in preference to export to the deep sea, organic carbon produced during the NEM is retained within the surface layer in the dissolved form, the magnitude of which is of the same order as the POC export during the entire year. This DOC pool is expected to be labile and may serve as a nutrient source for the microbial loop in the surface layer during the following relatively oligotrophic SI season (Azam et al., 1994; Madhupratap et al., 1996b, 2001). Circumstantial evidence that lends support to this hypothesis includes higher biomasses of heterotrophic bacteria, nano- and microheterotrophs as well as mesozooplankton during the SI. As previously mentioned, the biomass of zooplankton >2 mm is maximal during the SI period most likely in response to greater food supply. If the zooplankton in this region met their food demand primarily through grazing on microzooplankton (Smith et al.,

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1998), then the nano- and microzooplankton stocks in turn should be sustained by the availability of either heterotrophic bacteria (growing on labile DOC) or pico-plankton that make use of the regenerated nutrients. Significantly, the abundances of both pico-autotrophs (primarily Prochlorococcus) and heterotrophic bacteria are close to their maxima during the SI, and this perhaps ensures enough carbon supply for other constituents of the microbial food chain. 6.6.3

Role of Arabian Sea as a source or sink of carbon dioxide (CO2)

Areas experiencing upwelling and convective mixing can serve either as net sources or net sinks of CO2 depending upon the balance between gas emission to the atmosphere, due to warming of cold waters derived from depth, and biological uptake for photosynthesis, stimulated by the supply of nutrients. In the Arabian Sea, the physical effect appears to overwhelm the biological one. Coastal upwelling and winter mixing have been found to elevate partial pressure of carbon dioxide (pCO2) in surface waters up to ~700 and 450 μatm. respectively (Körtzinger et al., 1997; Goyet et al., 1998; Sarma, 1998). While peak values during the SWM are among the highest recorded at the sea surface, the areal extent of elevated pCO2 is more during the NEM. Even during the intermonsoon periods, surface-water pCO2 does not fall below the atmospheric value (George et al., 1994; Sarma et al., 1998). This supersaturation drives a flux of CO2, perennially and all across the basin with marked spatial and temporal variability, from the ocean to the atmosphere; its magnitude has been estimated to be 7–79 Tg y−1 (Goyet et al., 1998; Sarma et al., 1998). Given the uncertainties involved in this estimate and also of the above-mentioned export to the deep sea, the magnitudes of CO2 efflux to the atmosphere and POC export to the deep sea are quite comparable. However, it must be appreciated that the scales on which these processes operate are quite different. Thus, while the Arabian Sea serves as a net source of CO2 on the short-term and regional scales, it must act as a sink on long-term basis since a significant amount of carbon exported to the deep sea is preserved in the sediments (Smith, 2001).

6.7

Oxygen-deficient zones

For reasons given earlier, dissolved O2 concentrations in much of the mesopelagic Arabian Sea (depth range ~150–1000 m) are below 0.1 ml l−1 (~4 μM, Fig. 6.14a), making it one of the largest and the most intense O2-deficient zones (ODZs) in the world. The most O2-deficient waters are found north of ~12° N latitude (Fig. 6.14a), the approximate position of the zero wind stress curl (Warren, 1994), where suboxic (O2 <1 μM) conditions profoundly influence

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Fig. 6.14 Distributions of (a) O2 (μM) and (b) NO2− (μM) in the upper kilometre along the track of TN039 cruise of US JGOFS. The inset in panel (b) shows the station locations and the geographical extent of the Arabian Sea denitrification zone as demarcated by the 0.5 μM NO2− contour (Naqvi, 1991).

biogeochemical processes in the mesopelagic zone and over the continental margins.

6.7.1

Denitrification

The near absence of O2 results in vigorous denitrification [bacterial conversion of NO3− to molecular nitrogen (N2)] within a layer that is several hundred metres thick and spreads over an area measuring around 1.4 × 106 km2 (Naqvi, 1994). This zone can be easily identified by the accumulation of nitrite (NO2−), the first intermediate of the reduction sequence (NO3− → NO2− → NO → N2O → N2) (Figs 6.14b and 6.15a). Note that NO2− is also produced through assimilatory reduction of NO3− by phytoplankton and the oxidation of NH4+ (nitrification), but both these processes occur in oxic waters close to the surface (Codispoti & Christensen, 1985). While the primary NO2− maximum resulting from nitrification/assimilatory NO3− reduction is ubiquitously found near the base

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of the surface mixed layer, the secondary NO2− maximum (SNM) associated with denitrification is confined only to a few areas of the open ocean (besides the Arabian Sea the other such areas are the eastern-boundary upwelling zones off Peru and Mexico in the Pacific and off Namibia in the Atlantic – Hattori, 1983). What distinguished the Arabian Sea from the other regions is that here the most vigorous denitrification occurs offshore, well away from the intense upwelling sites in the western Arabian Sea (Fig. 6.14b-inset). This is because for appreciable denitrification to occur, as inferred from the accumulation of secondary NO2−, O2 concentration must fall below an abruptly defined threshold value of ~1μM (~0.02mll−1, Fig. 6.15a). This threshold appears to represent an ecosystem switch that brings about major alterations in biogeochemical cycles of other polyvalent elements as well, as we shall see later. Outflows from the Red Sea and the Persian Gulf and advection of waters from the south in the western Arabian Sea generally keep minimum O2 levels marginally above this threshold (note the elevated O2 concentrations associated with the Persian Gulf Outflow near the northern end of the section in Fig. 6.14a). A seasonal undercurrent bringing relatively oxygenated waters from the south during the SWM similarly suppresses denitrification just off the Indian continental margin (Naqvi et al., 1990). The significance of this ecosystem switch must be appreciated in the context of the huge volume of water that presently has dissolved O2 in the non-reducing concentration range of 1–5 μM (Fig. 6.14). It not only accounts for the observed spatial and temporal (seasonal as well as inter-annual) changes in denitrification (Naqvi et al., 1990; Morrison et al., 1998), but also implies that even minute human-induced perturbations of oxygen and carbon budgets have the potential to profoundly impact biological diversity and chemical fluxes, making the Arabian Sea a sensitive barometer of global change (Mantoura et al., 1993). Geological records provide evidence for such shifts in the past arising from, and most likely contributing to, climatic cycles (Altabet et al., 2002). The loss of NO3− through denitrification produces a minimum within the ODZ where its concentration is lower by 8–12 μg-atm. N l−1 than the amounts expected from the relationships between NO3− and temperature/salinity/O2/ PO4−3 in the more oxygenated waters (Naqvi, 1994 and references therein; Bange et al., 2000; Codispoti et al., 2001). Measurements of N2/Ar ratio have led to estimates of excess N2 (maximum ~20 μg-atm. N l−1) which substantially exceed NO3− deficits (Codispoti et al., 2001). However, even the seemingly conservative estimates of NO3− deficits have yielded rates of water-column denitrification ranging from 21 to 33 Tg N y−1 (Naqvi, 1987; Howell et al., 1997; Bange et al., 2000), very close to the estimate based on the activity of respiratory electron transport system (ETS) (Naqvi & Shailaja, 1993). Thus, the Arabian Sea’s contribution to total oceanic pelagic denitrification (80–150 Tg N y−1, Gruber & Sarmiento, 1997; Codispoti et al., 2001) is very substantial and globally significant.

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Within the ODZ, nitrous oxide (N2O), another major intermediate of denitrification (and a byproduct of nitrification), shows a trend of variability quite different from that of NO2− (Fig. 6.15), but similar to that observed in the ODZs of the Pacific Ocean (Codispoti & Christensen, 1985). That is, N2O concentration generally increases non-linearly with the depletion in O2 until the environment turns reducing; thereafter, concomitant with the accumulation of secondary NO2, a rapid fall in N2O concentration takes place. Accordingly, the SNM is characterized by a minimum in N2O concentration (<10 nM), whereas the oxic–suboxic interfaces are characterized by peak N2O levels exceeding 50 nM (Law & Owens, 1990; Naqvi & Noronha, 1991). Attempts have been made to evaluate the relative importance of nitrification, denitrification and coupling between the two processes as pathways for N2O production by

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investigating the natural abundance of N and O isotopes in N2O (Yoshinari et al., 1997; Naqvi et al., 1998). The complex patterns observed suggest the likelihood of involvement of several formative mechanisms. The upper N2O maximum is located just below the mixed layer, and so it sustains large inputs of N2O to surface waters, in turn supporting unusually high atmospheric flux from this region (Bange et al., 2001 and references therein). Moreover, the close proximity of intense sources and sinks results in a very rapid turnover of N2O (Naqvi & Noronha, 1991). In addition to the perennial open-ocean suboxic zone, as previously mentioned, O2-deficient conditions also develop seasonally in shallow subsurface waters over the western continental shelf of India (Fig. 6.16). The O2-deficiency is most severe in the autumn when the entire shelf is covered by waters with O2 < 0.5 ml l−1 (22 μM). The area (~200 000 km2) of this zone is an order of magnitude more than the area of the largest human-induced ODZ in the Gulf of Mexico. Although, as described earlier, the causative factors for this phenomenon are primarily natural and have been recognized for quite some time (Banse, 1959; Carruthers et al., 1959), there is compelling evidence pointing to an intensification of this system in recent years (Naqvi et al., 2000). For example, episodes of complete denitrification and the ensuing sulphate reduction (Fig. 6.16), extremely rare in open coastal waters, have never been reported previously from the Arabian Sea. It is likely that fertilizer inputs from land have further enhanced the naturally high PP rates in coastal waters bringing about an ecosystem shift in recent years, but subtle changes in hydrography cannot be excluded as an additional or alternative cause. In spite of the seasonal existence and small area of the coastal ODZ, its contribution to the overall denitrification rate appears to be quite significant (~5 Tg N y−1). Biogeochemical transformations involving N2O in the coastal and open ocean ODZs appear to be different in that the coastal suboxic zone often experiences accumulation of N2O to levels (up to ~0.5 μM) that have not been observed anywhere else in the ocean (Fig. 6.16). The high NO2− and low NO3− concentrations in these waters strongly indicate the prevalence of reducing conditions. Transient build-up of N2O at μM level during denitrification was also observed during deck incubation of initially oxic samples (Naqvi etal., 2000). These results indicate that in shallow, rapidly denitrifying systems subjected to frequent aeration due to turbulence, the activity of N2O-reductase may be suppressed as a result of which a large fraction of nitrogen undergoing bacterial reduction ends up as N2O (Naqvi et al., 2000). Further, due to the proximity of N2O-rich waters to the sea surface, the N2O so produced can easily escape to the atmosphere. It has been estimated that the annual flux of N2O to the atmosphere from the Arabian Sea is in the range 0.36–1.09 Tg N2O (Naqvi et al., 2000; Bange et al., 2001), with the upper limit amounting to as much as one-fifth of the total oceanic emissions of N2O (Codispoti et al., 2001). Elevated emissions of methane to the atmosphere have also been reported from the

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Stations E 102 103

101

100

099

098

0

097

W

50 100

Temperature (°C)

0 50 100

Salinity

0 50 100

O2 (μM)

0 0 Depth (m)

50 NO2– (μM)

100 0 50

NO3– (μM)

100 0 405

0

50 100

N2O (nM)

0 7 50 NH4+ (μM)

100 0

19 50 100

0 H2S (μM) 100 150 50 Distance from coast (km)

Fig. 6.16 Distributions of temperature, salinity, oxygen, inorganic nitrogen species and hydrogen sulphide along a transect of the continental shelf off Bombay during 6–7 October 1999. Reproduced from Naqvi et al. (2000).

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Arabian Sea, but unlike N2O their impact on global methane budget is insignificant (Owens et al., 1991). 6.7.2

Intermediate nepheloid layer

Beam attenuation profiles in the Arabian Sea often show the presence of a secondary particle maximum, or intermediate nepheloid layer (INL), between 150 and 350m (Naqvi et al., 1993; Gundersen et al., 1998; Morrison et al., 1999) identical to those found in the ODZs of the eastern tropical Pacific Ocean (Spinrad et al., 1989 and references therein). The INL shows a remarkable association with the SNM; the two features also undergo very similar offshore intensification (Naqvi et al., 1993; Morrison et al., 1999). Moreover, maxima in bacterial abundance (Fig. 6.17) particulate protein concentration and ETS activity (a measure of respiration rate) have also been found to be associated with the INL. These results strongly point to an in situ biological origin of the INL (bacterial proliferation) rather than an advective one (offshore extension of the bottom nepheloid layer). Intriguingly, the organic carbon demand within the SNM, derived from the ETS activity (Naqvi et al., 1993) or bacterial production (Ducklow, 1993), is about an order of magnitude more than the sinking flux of POC from the surface layer as measured by drifting sediment traps or estimated from global relationships between the PP and POC export (Naqvi & Shailaja, 1993; Naqvi et al., 1993). What produces the bacterial biomass maximum in denitrifying waters – whether it is due to lower mortality

BAC (m–1) 0.35

0.43

0.49

Bact. abundance (108 cells cm–3) 10 20

Nitrite (μM) 1

3

5

0

Depth (km)

0.2 0.4 0.6 0.8 (a)

(b)

(c)

1.0 Fig. 6.17 Typical profiles of (a) beam attenuation coefficient (BAC, m−1, from Naqvi et al., 1993) and (b) bacterial abundance (108 cells l−1, from Ducklow, 1993) within the denitrifying zone. The sampling sites, located close to each other (around 15° N, 67° E) had similar NO2− distributions (μM) (c), open circles corresponding to microbial data. Subsequent measurements (Bess Ward, Amal Jayakumar & Wajih Naqvi, unpublished) have confirmed the co-occurrence of these features.

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of heterotrophic bacteria, as a result of exclusion of microscopic grazers from the suboxic zone, or due to enhanced growth under suboxic conditions – is not clear. The former mechanism requires the critical O2 concentration for the exclusion of micrograzers to be exactly the same as the threshold value for denitrification, which is possible but highly unlikely. As for the apparent inadequacy of the sinking flux to meet the observed respiration rates, additional modes of organic carbon input to the denitrifying layer, besides sedimentation of particles from the surface, should be looked for. One possibility is that the organic matter is transported either quasi-horizontally from the continental margins (Somayajulu et al., 1996) or vertically through diurnal migrations of organisms particularly the myctophids, which are especially adapted to tolerate suboxic conditions (see below) and are amazingly abundant in the Arabian Sea (Banse, 1994; Morrison et al., 1999). The other explanation could be that the denitrifiers are able to efficiently utilize DOC (Naqvi et al., 1993) or transparent exopolymer particles (Kumar et al., 1998). If the heterotrophic carbon utilization is dependent on the redox status of the environment, then it has important implications for biogeochemical cycling. 6.7.3

Other redox-sensitive elements

Reducing conditions prevailing within the core of the ODZ also affect the cycling of some other polyvalent elements such as manganese (Mn), iron (Fe) and iodine (I). For example, dissolved Mn (d-Mn) concentrations rise abruptly when O2 levels approach ~1 μM, exhibiting a pattern very similar to that of NO2− (Fig. 6.15c). Evidently, this increase is caused by in situ reductive solubilization of particulate Mn oxyhydroxides (Lewis & Luther, 2000). Besides this maximum (d-Mn up to 8nM against the surface and deep water values of 1–2nM and ≤0.6 nM, respectively), which coincides with the SNM, another welldeveloped maximum (d-Mn up to 6 nM) is also found at ~600 m depth in the northern Arabian Sea. The deeper maximum is believed to arise from horizontal advection of d-Mn-enriched layers formed as a result of diffusion of Mn(II) from the reducing continental margin sediments (Saager et al., 1989; Lewis & Luther, 2000). An almost similar distribution pattern has been reported for d-Fe (Saager et al., 1989). Another species that undergoes reduction in the Arabian Sea is iodate (IO3−). Subsurface maxima in iodide (I−) coincident with minima in IO3− within the ODZ have been found even outside the zone of active denitrification, prompting the suggestion that reduction of IO3− might precede that of NO3− (the free energy change associated with the oxidation of organic matter with both oxidants is very similar; Farrenkopf et al., 1997). Thus, in addition to NO3− , there are a number of other important electron acceptors [Mn(III and IV), Fe(III) and I(V)] in the non-sulphidic but O2-poor environments of the Arabian Sea – both in the water and in sediments. It is quite likely that most of them, if not all, are used by the same facultative bacteria (e.g. Shewenella putrefaciens

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strain MR-4; Farrenkopf et al., 1997). Some of these also have the potential to directly react with nitrogen species [e.g. Mn(IV) with NH4+ , and Mn(II) with NO3− ] to produce N2 (Luther et al., 1997). Such interactions are expected to be particularly important in the coastal suboxic zone where the proximity to sediments and changes in redox state of overlying waters are expected to result in a rapid turnover of redox-sensitive elements. Also, the greater mobilization in suboxic waters of the redox-sensitive Fe, a well-known limiting micronutrient for plankton growth, probably supplements the eolian supply of Fe to the region and ensures its adequate availability for plankton growth (Measures & Vink, 1999; Smith, 2001). 6.7.4

Biological effects

The intense O2-deficiency strongly influences the abundance and composition of marine organisms in the mesopelagic realm. As in the eastern tropical North Pacific (ETNP), most zooplankton are excluded from the ODZ due to their inability to cope with suboxic conditions (Wishner et al., 1998). The O2 deficiency also affects the migration of the majority of marine organisms. For instance, the fraction of the total biomass (0–1000 m) that migrates into the upper layer (0–100 m) has been found to be the lowest at stations with most pronounced O2 minimum (Wishner et al., 1998). However, some species of diel migrators, which are apparently well adapted to tolerate low O2 levels, are abundant within the suboxic zone, with the larger size group (>2 mm) of such organisms (euphausiids and myctophids) reaching depths of up to 400 m (Wishner et al., 1998). Such diel migration of organisms is clearly seen in the backscatter data (Fig. 6.18) taken with the acoustic Doppler current profiler (ADCP). Copepods were found to undergo little migration at this location (Smith et al., 1998), and the majority of organisms detected by the ADCP are believed to be myctophids (Morrison et al., 1999) whose biomass in the Arabian

Fig. 6.18 Acoustic backscatter for 24 hours around 19° N, 67° E indicating diurnal migration of organisms (mostly myctophids). Reproduced from Morrison et al. (1999) (see Color Plate 4).

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Sea has been estimated to be of the order of 100 million tonnes (GLOBEC, 1993). The lower suboxic/oxic interface (O2 ~0.05–0.1 ml l−1) is distinguished by an increase in zooplankton biomass (Fig. 6.19), a feature that is probably common to all regions containing pronounced ODZs (Wishner et al., 1998). This layer, about 100 m thick, is also characterized by elevated zooplankton feeding rates and POC concentration as well as particle aggregation (Morrison et al., 1999). Apparently, the food chains within this zone are short and the biological modification of sinking POC far more active than in the overlying suboxic water column (Wishner et al., 1998). The calanoid copepod Lucicutia grandis is a prominent indicator species of this environment (Gowing & Wishner, 1998). Reproducing throughout the year, the omnivorous L. grandis consumes a variety of materials (surface derived, deep sea detritus, zooplankton and aggregates; Wishner et al., 2000). Patterns of vertical distribution of various taxa contributing to the subsurface biomass peak vary depending upon their

Oxygen (ml/l) Carbon (m M Cm–3) 0.0

0.1

0.2

OtherTaxa (#/1000 m3)

Copepods (#/1000 m3) 0

20,000

40,000

0

500

1,000

1,500

300 x

400

Oxygen Carbon

x Ostracods Salps/Doliolids Cyclothone

Total Copepods Calanoids Non-calanoids

Pressure (dbar)

500 x 600

700

800 x 900 x 1000

Fig. 6.19 Profiles of O2 and mesozooplankton components from a vertically stratified MOCNESS night tow at US JGOFS Sta. S7 on 17 December 1995. Reproduced from Morrison et al. (1999).

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ability to tolerate low O2 concentrations: gelatinous taxa such as salps and doliolids occur in more O2-depleted shallower waters than do crustaceans (copepods and shrimps) and fish (Cyclothone). Overall, the structure and function of the mid-water ecosystem seem to exhibit a remarkable long-term stability as indicated by small seasonal changes at any given site, except in the nearshore waters, showing the strong influence of O2-deficient conditions on biology of the deep Arabian Sea (Smith et al., 2002). A secondary fluorescence maximum (SFM), almost exclusively caused by Prochlorococcus, is associated with the upper oxic/hypoxic interface (Johnson et al., 1999). This strain of Prochlorococcus, possibly different from the one found in the oligotrophic surface waters, is well-adapted to low light and O2 levels. Similar SFM has also been reported from the ETNP, and is believed to be related to the presence of strong ODZ.

6.8

Benthic processes

Chemical and biological processes in the sediments and benthic boundary layer (BBL) are important contributors to oceanic biogeochemical cycles, especially in the Arabian Sea due to its uncommon geographical setting. The oceanographic conditions experienced by various margins (e.g. Somalia/Oman versus India/Pakistan) are widely different, which in conjunction with the extensive mid-depth O2 deficiency produce a variety of BBL and sedimentary environments with respect to, among other factors, food supply, redox status and the nature and activity of benthic communities (Cowie, 2002). The most striking feature of the composition of Arabian Sea sediments is the unusually high organic carbon accumulation – among the highest in the world – occurring along the continental margins (Paropkari et al., 1992 and references therein). There has been considerable debate as to whether this accumulation is primarily due to enhanced carbon preservation favoured by low bottom-water O2 levels (Paropkari et al., 1992) or high biological productivity (Pedersen et al., 1992; Calvert et al., 1995). Those who do not support the preservation hypothesis argue that the sedimentary organic carbon content peaks in the mid-slope region, substantially below the core of the O2 minimum, and the differences in the composition of organic matter within and outside the ODZ off Oman (Pedersen et al., 1992), India (Calvert et al., 1995) and Pakistan (Cowie et al., 1999) are usually small. Consequently, other factors such as winnowing and offshore changes in productivity, depth (including bottom topography), sedimentation rate and benthic fauna may exert important control on both the concentration and composition of sedimentary organic carbon (Cowie, 2002). The suboxic/anoxic conditions prevailing in marginal sediments of the Arabian Sea appear to affect, to a very substantial extent, cycling and global

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budgets of phosphorus, nitrogen and sulphur and of redox sensitive metals (Fe and Mn). Burial of authigenic P in reducing margin sediments can alter oceanic P inventory and biological productivity (Ganeshram et al., 2002), whereas sedimentary denitrification leads to significant production of N2 and N2O (Cowie, 2002). Limited measurements of sedimentary denitrification over the Indian continental shelf have yielded rates (80–667 μmol m−2 d−1 corresponding to 0.4–3.5 Tg N y−1 for all Arabian Sea shelves) that are of the same order as those reported from other areas (Naik & Naqvi, 2002). However, as compared to other upwelling environments (e.g. off Peru), sulphate reduction rates in the margin sediments of the Arabian Sea are modest (e.g. 20–190 nmol cm−3 d−1 in the top 10 cm along the Pakistan margin; Schmaljohann et al., 2001) and the porewater sulphate concentrations are not greatly depleted within the upper 50 cm (Passier et al., 1997; Cowie etal., 1999). These results reflect the refractory nature of organic matter or an inhibition of its post-oxic decay (Cowie, 2002). The concentration of free sulphide is quite low which could to a large extent be due to the scavenging by Fe to form pyrite (Passier et al., 1997; Schmaljohann et al., 2001). Mats of filamentous bacteria (mostly Thioploca but possibly also Beggiatoa), that facilitate oxidation of reduced sulphur by NO3− , have been reported from the continental slope off Oman (Levin et al., 1997) as well as Pakistan (Schmaljohann et al., 2001). However, these unusual chemoautotrophs can account for only a part of the substantial CO2 fixation occurring in the upper 5 cm of the sedimentary column as well as of the total sulphide removal (Schmaljohann et al., 2001). The benthic biology is greatly impacted where the ODZ impinges on the seafloor (Gage et al., 2000 and papers therein). In the northeastern Arabian Sea where the ODZ is particularly intense, the lack of macrofauna activity (bioturbation) is manifested by the accumulation of laminated (varved) sediments that are ideally suited for the paleoceanographic reconstruction at an exceptionally fine resolution (von Rad et al., 1996). Interestingly, while the diversity and evenness of benthic fauna is reduced, the total biomass may be higher within the sediments in contact with the ODZ than those exposed to more oxygenated waters (Levin et al., 2000). The benthic community within the ODZ is largely composed of surface-feeding polychaetes, whereas mollusks and most crustaceans are common below this zone (Levin et al., 2000). Some of the polychaetes have been found to undergo morphological modifications in order to adapt to low O2 levels by increasing surface area of their gas-exchange organs (size and number of branchiae; Lamont & Gage, 2000). As the O2 levels rise above the threshold of 0.15–0.20 ml l−1, the macrofauna take advantage of the abundant food supply, resulting from reduced utilization of the downward flux of organic matter within the ODZ, producing an abundance maximum (at 700–850m off Oman; Levin et al., 2000). In the abyssal region, the onshore–offshore gradients in various parameters of benthic processes generally reflect those in PP and related vertical fluxes of

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POC (Table 6.3). The most eutrophic region of the western Arabian Sea is characterized by enhanced pigment concentrations in sediments as well as higher biomasses of several groups of organisms, while the lowest values of these parameters have been recorded in the relatively oligotrophic southern part (Pfannkuche & Lochte, 2000 and references therein). For example, the biomass of metazoan macrofauna, comprising mostly crustaceans and polychaetes, at the site of the southern Arabian Sea trap (SAST) is about one-fifth of the biomass at the location of the western Arabian Sea trap (WAST) (Witte, 2000). The metazoan meiofauna exhibits a similar distribution pattern: the biomass of the nematodes, accounting for most of the meiofauna, being 13.3 and 2.6mgCm−2 at WAST and SAST sites, respectively (Sommer & Pfannkuche, 2000). However, what is surprising is that while the regional gradients in the standing stocks of both macro- and meiobenthos reflect food availability within the Arabian Sea, the biomasses for the region as a whole are not much higher than those in other areas (e.g. NE Atlantic) where organic carbon fluxes to the deep sea are smaller (Sommer & Pfannkuche, 2000; Witte, 2000). On the other hand, the bacterial activity and biomass do reflect high organic carbon inputs to the deep sea sediments, being substantially higher in the Arabian Sea than in other regions, and also exhibiting similar regional variability to the POC flux (Boetius et al., 2000). Moreover, the microbial parameters exhibit appreciable seasonal changes although these are less pronounced than the geographical changes. The share of bacteria thus appears to be higher – 30–70% of the POC turnover (Boetius et al., 2000) – relative to other structural components of the benthic community (Witte, 2000).

Table 6.3 Selected biogeochemical characteristics of sediments at three sites of Indo-German sediment traps (WAST – Western Arabian Sea Trap; CAST – Central Arabian Sea Trap; SAST – Southern Arabian Sea Trap) – after Pfannkuche and Lochte (2000)

PP (g C m−2 y−1) POC flux (1 km) (g C m−2 y−1) Chloroplastic pigments (μg cm−3 surface sediment) Sediment community oxygen consumption (mmol O2 m−2 d−1) Benthic carbon remineralisation (g C m−2 y−1) Bacterial secondary production (0–5 cm) (mg C m−2 d−1) Bacterial biomass (0–5 cm) (g C m−2) Meiofauna abundance (Ind./10 cm2) Living foraminifera biomass (mg C m−2) Macrofauna biomass (mg C m−2)

WAST 16° N, 60° E, 4030 m

CAST 15° N, 65° E, 3960 m

SAST 10° N, 65° E, 4410 m

244 3.2 5.2

204 1.9 2.1

154 1.2 0.9

5.1

3.3

1.1

13.5 0.55

10.1 0.20

3.1 0.07

0.8 204 250 106

0.6 227 31 23

0.4 125 11 21

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Conforming to the results of microbial studies, rates of sediment community oxygen consumption (SCOC), measured in situ through deployment of bottom landers, are also some of the highest (0.9–6.3 mmol O2 m−2 d−1) recorded in the deep ocean. As expected, maximal SCOC values occur in the western and northern sectors, decreasing towards the central and southern parts (Witte & Pfannkuche, 2000). As in case of microbial parameters, there is some evidence that SCOC responds to seasonal changes in the overhead PP. However, the sinking POC flux alone seems to be insufficient to support benthic carbon remineralization, especially at the WAST location, requiring additional modes of POC supply such as accelerated deposition of labile organic matter facilitated by the eolian supply of dust particles (the mineral ballast), enhanced rain efficiencies and lateral advection (Witte & Pfannkuche, 2000). The larger supply of organic matter to the sediments of the western Arabian Sea leads to the development of anoxic conditions a few centimetres below the sediment water interface, a phenomenon not noticed in the southern region (Boetius et al., 2000). However, as a cautionary note, it must be pointed out that the geographical and temporal coverage of the Arabian Sea for benthic biogeochemical studies has been severely inadequate so far, since not all national JGOFS programmes included benthic studies. Future work will hopefully remedy this shortcoming to achieve a comprehensive understanding of the biogeochemistry of this unique region. References Altabet, M.A., Higginson, M.J. and Murray, D.W. (2002) The effect of millennial-scale changes in Arabian Sea denitrification on atmospheric CO2. Nature, 415, 159–162. Angel, M. (ed.) (1984) Marine Science of the North-west Indian Ocean and Adjacent Waters, Deep-Sea Research, vol. 31, no. 6–8, Pergamon Press, London. Azam, F., Fenchel, T., Field, J.G., Gray, J.S., Meyer-Reil, L.A. and Thingstad, F. (1983) The ecological role of water-column microbes in the sea. Marine Ecology Progress Series, 10, 257–263. Azam, F., Steward, G.F., Smith, D.C. and Ducklow, H.W. (1994) Significance of bacteria in carbon fluxes in the Arabian Sea. Proceedings of the Indian Academy of Sciences (Earth and Planetary Sciences), 103, 341–351. Baars, M.A. (1994) Monsoons and Pelagic Systems, National Museum of Natural History, Leiden. Baars, M. and Brummer, G.J. (1995) Workshop reviews results from Netherlands Indian Ocean Programme. U.S. JGOFS News, 6, 10–11. Baars, M.A. and Oosterhuis, S.S. (1997) Zooplankton biomass in the upper 200 m in and outside the seasonal upwelling areas of the western Arabian Sea, in Pelagic Biogeography ICoPBII Proceedings of the Second International Conference (eds A.C. Pierrot-Bults and S. van der Spoel), IOC/UNESCO, Paris, pp. 39–52. Bange, H.W., Andreae, M.O., Lal, S., Law, C.S., Naqvi, S.W.A., Patra, P.K., Rixen, T. and UpstillGoddard, R.C. (2001) Nitrous oxide emissions from the Arabian Sea: a synthesis. Atmospheric Chemistry and Physics, 1, 61–71. Bange, H.W., Rixen, T., Johansen, A.M., Siefert, R.L., Ramesh, R., Ittekkot, V., Hoffmann, M.R. and Andreae, M.O. (2000) A revised nitrogen budget for the Arabian Sea. Global Biogeochemical Cycles, 14, 1283–1297.

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Banse, K. (1959) On upwelling and bottom-trawling off the southwest coast of India. Journal of the Marine Biological Association of India, 1, 33–49. Banse, K. (1987) Seasonality of phytoplankton chlorophyll in the central and northern Arabian Sea. Deep-Sea Research, 34, 713–723. Banse, K. (1994) On the coupling of hydrography, phytoplankton, zooplankton, and settling organic particles offshore in the Arabian Sea. Proceedings of the Indian Academy of Sciences (Earth and Planetary Sciences), 103, 125–161. Banse, K. and English, D.C. (2000) Geographical differences in seasonality of CZCS-derived phytoplankton pigment in the Arabian Sea for 1978–1986. Deep-Sea Research II, 47, 1623–1677. Banse, K. and McClain, C.R. (1986) Winter blooms of phytoplankton in the Arabian Sea as observed by the coastal zone color scanner. Marine Ecology Progress Series, 34, 201–211. Barber, R.T., Marra, J., Bidigare, R.R., Codispoti, L.A., Halpern, D., Johnson, Z., Latasa, M., Goericke, R. and Smith, S.L. (2001) Primary productivity and its regulation in the Arabian Sea during 1995. Deep-Sea Research II, 48, 1127–1172. Barlow, R.G., Mantoura, R.F.C. and Cummings, D.G. (1999) Monsoonal influence on the distribution of phytoplankton pigments in the Arabian Sea. Deep-Sea Research II, 46, 677–689. Bauer, S., Hitchcock, G.L. and Olson, D.B. (1991) Influence of monsoonally forced Ekman dynamics upon surface layer depth and plankton biomass distribution in the Arabian Sea. Deep-Sea Research, 38, 531–553. Boetius, A., Ferdelman, T. and Lochte, K. (2000) Bacterial activity in sediments of the deep Arabian Sea in relation to vertical flux. Deep-Sea Research II, 47, 2835–2875. Bottger-Schnack, R. (1996) Vertical structure of small metazoan plankton, especially non-calanoid copepods. 1. Deep Arabian Sea. Journal of Plankton Research, 18, 1073–1101. Brandes, J.A., Devol, A.H., Jayakumar, D.A., Yoshinari, T. and Naqvi, S.W.A. (1998) Isotopic composition of nitrate in the central Arabian Sea and eastern tropical North Pacific: A tracer for mixing and nitrogen cycles. Limnology and Oceanography, 43, 1680–1689. Brock, J.C., Sathyendranath, S. and Platt, T. (1998) Biohydro-optical classification of the northwestern Indian Ocean. Marine Ecology Progress Series, 165, 1–15. Broerse, A.T.C., Brummer, G.J.A. and Van Hinte, J.E. (2000) Coccolithophore export production in response to monsoonal upwelling off Somalia (northwestern Indian Ocean). Deep-Sea Research II, 47, 2179–2205. Brown, S.L., Landry, M.R., Barber, R.T., Campbell, L., Garrison, D.L. and Gowing, M.M. (1999) Picophytoplankton dynamics and production in the Arabian Sea during the 1995 southwest monsoon. Deep-Sea Research II, 46, 1745–1768. Buesseler, K.O., Ball, L., Andrews, J., Benitez-Nelson, C., Belastock, R., Chai, F. and Chao, Y. (1998) Upper ocean export of particulate organic carbon in the Arabian Sea derived from thorium-234. Deep-Sea Research II, 45, 2461–2487. Burkill, P.H. (ed.) (1999) ARABESQUE: UK JGOFS Process Studies in the Arabian Sea, Deep-Sea Research II, vol. 46, no. 3–4, Pergamon Press, London. Burkill, P.H., Leakey, R.J.G., Owens, N.J.P. and Mantoura, R. (1993) Synechococcus and its importance to the microbial foodweb of the northwestern Indian Ocean. Deep-Sea Research II, 40, 773–782. Calvert, S.E., Pedersen, T.F., Naidu, P.D. and von Stackelberg, U. (1995) On the organic carbon maximum on the continental slope of the eastern Arabian Sea. Journal of Marine Research, 53, 269–296. Campbell, L., Landry, M.R., Constantinou, J., Nolla, H.A., Brown, S.L., Liu, H. and Caron, D.A. (1998) Response of microbial community structure to environmental forcing in the Arabian Sea. Deep-Sea Research II, 45, 2301–2325. Capone, D.G., Subramaniam, A., Montoya, J.P., Voss, M., Humborg, C., Johansen, A.M., Siefert, R.L. and Carpenter, E.J. (1998) An extensive bloom of the N2-fixing cyanobacterium Trichodesmium erythraeum in the central Arabian Sea. Marine Ecology Progress Series, 172, 281–292. Caron, D.A. and Dennett, M.R. (1999) Phytoplankton growth and mortality during the 1995 northeast monsoon and spring intermonsoon in the Arabian Sea. Deep-Sea Research II, 46, 1665–1690.

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Smith, S.L. (ed.) (2000) The 1994–1996 Arabian Sea Expedition: Oceanic Response to Monsoonal Forcing, Part 3, Deep-Sea Research II, vol. 47, no. 7–8, Pergamon Press, London. Smith, S.L. (ed.) (2001) The 1994–1996 Arabian Sea Expedition: Oceanic Response to Monsoonal Forcing, Part 4, Deep-Sea Research II, vol. 48, no. 6–7, Pergamon Press, London. Smith, S.L. and Codispoti, L.A. (1980) Southwest monsoon of 1979: Chemical and biological response of Somali coastal waters. Science, 209, 597–600. Smith, S.L., Morrison, J.M. and Codispoti, L.A. (2002) The role of the oxygen minimum zone in biogeochemical cycles, in Report of the Indian Ocean Synthesis Group on the Arabian Sea Process Study, JGOFS Report No. 35 (eds L. Watts, P. Burkill and S. Smith), Scientific Committee on Oceanic Research, International Council of Scientific Unions, Bergen, pp. 57–64. Smith, S., Roman, M., Prusova, I., Wishner, K., Gowing, M., Codispoti, L., Barber, R., Marra, J. and Flagg, C. (1998) Seasonal response of zooplankton to monsoonal reversals in the Arabian Sea. Deep-Sea Research II, 45, 2369–2403. Somayajulu, B.L.K., Sarin, M.M. and Ramesh, R. (1996) Denitrification in the eastern Arabian Sea: Evaluation of the role of continental margins using Ra isotopes. Deep-Sea Research II, 43, 111–117. Sommer, S. and Pfannkuche, O. (2000) Metazoan meiofauna of the deep Arabian Sea: Standing stocks, size spectra and regional variability in relation to monsoon induced enhanced sedimentation regimes of particulate organic matter. Deep-Sea Research II, 47, 2957–2977. Spinrad, R.W., Glover, H., Ward, B.B., Codispoti, L.A. and Kullenberg, G. (1989) Suspended particle and bacterial maxima in Peruvian coastal waters during a cold water anamaly. Deep-Sea Research, 36, 715–733. Stelfox, C., Burkill, P., Edwards, E., Harris, R. and Sleigh, M. (1999) The structure of zooplankton communities, in the 2 to 2000 μm size range, in the Arabian Sea during and after the SW monsoon, 1994. Deep-Sea Research II, 46, 815–842. Suthhof, A., Jennerjahn, T., Rixen, T., Schäfer, P., Tiemann, J., Haake, B. and Ittekkot, V. (1999) Time-varying fluxes and burial of organic carbon along the continental margins in the northeastern Arabian Sea, in Land-Sea Link in Asia, Proceedings of International Workshop on Sediment Transport and Storage in Coastal Sea-ocean Systems (eds Y. Saito, K. Ikehara and H. Katayama), JAMSTEC and Geological Survey of Japan, Tsukuba, pp. 275–280. Swallow, J. (1984) Some aspects of the physical oceanography of the Indian Ocean. Deep-Sea Research, 31, 639–650. Tarran, G.A., Burkill, P.H., Edwards, E.S. and Woodward, E.M.S. (1999) Phytoplankton community structure in the Arabian Sea during and after the SW monsoon, 1994. Deep-Sea Research II, 46, 655–676. van Couwelaar, M. (1997) Zooplankton and micronekton biomass off Somalia and in the southern Red Sea during the SW monsoon of 1992 and the NE monsoon of 1993. Deep-Sea Research II, 44, 1213–1234. Van Veering, T.C.E., Helder, W. and Schalk, P. (eds) (1997) Netherlands Indian Ocean Program 1992–1993: First Results, Deep-Sea Research II, vol. 44, no. 6–7, Pergamon Press, London. Veldhuis, M.J.W., Kraay, G.W., van Bleijswijk, J.D.L. and Baars, M.A. (1997) Seasonal and spatial variability in phytoplankton biomass, productivity and growth in the northwestern Indian Ocean: the southwest and northeast monsoons, 1992–1993. Deep-Sea Research I, 44, 425–449. Vinogradov, M. and Voronina, N.M. (1961) Influence of the oxygen deficit on the distribution of plankton in the Arabian Sea. Okeanologiya, 1, 670–678. Warren, B.A. (1994) Context of the suboxic layer in the Arabian Sea. Proceedings of the Indian Academy of Sciences (Earth and Planetary Sciences), 103, 301–314. Watts, L., Burkill, P. and Smith, S. (eds) (2002) Report of the Indian Ocean Synthesis Group on the Arabian Sea Process Study, JGOFS Report No. 35, Scientific Committee on Oceanic Research, International Council of Scientific Unions, Bergen. Watts, L.J. and Owens, N.J.P. (1999) Nitrogen assimilation and the f-ratio in the northwestern Indian Ocean during an intermonsoon period. Deep-Sea Research II, 46, 725–743.

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Watts, L.J., Sathyendranath, S., Caverhill, C., Maass, H., Platt, T. and Owens, N.J.P. (1999) Modelling new production in the northwest Indian Ocean region. Marine Ecology Progress Series, 183, 1–12. Wiebinga, C.J., Veldhuis, M.J.W. and DeBaar, H.J.W. (1997) Abundance and productivity of bacterioplankton in relation to seasonal upwelling in the northwest Indian Ocean. Deep-Sea Research I, 44, 451–476. Wiggert, J.D. (2002) Modelling physical-biogeochemical interactions, in Report of the Indian Ocean Synthesis Group on the Arabian Sea Process Study, JGOFS Report No. 35 (eds L. Watts, P. Burkill and S. Smith), Scientific Committee on Oceanic Research, International Council of Scientific Unions, Bergen, pp. 79–84. Wiggert, J.D., Jones, B.H., Dickey, T.D., Brink, K.H., Weller, R.A., Marra, J. and Codispoti, L.A. (2000) The northeast monsoon’s impact on mixing, phytoplankton biomass and nutrient cycling in the Arabian Sea. Deep-Sea Research II, 47, 1353–1385. Wishner, K.F., Gowing, M.M. and Gelfman, C. (1998) Mesozooplankton biomass in the upper 1000 m in the Arabian Sea: Overall seasonal and geographic patterns and relationship to oxygen gradients. Deep-Sea Research II, 45, 2405–2432. Wishner, K.F., Gowing, M.M. and Gelfman, C. (2000) Living in suboxia: Ecology of an Arabian Sea oxygen minimum zone copepod. Limnology and Oceanography, 45, 1576–1593. Witte, U. (2000) Vertical distribution of metazoan macrofauna within the sediment at four sites with contrasting food supply in the deep Arabian Sea. Deep-Sea Research II, 47, 2979–2997. Witte, U. and Pfannkuche, O. (2000) High rates of benthic carbon remineralisation in the abyssal Arabian Sea. Deep-Sea Research II, 47, 2785–2804. Woodward, E.M.S., Rees, A.P. and Stephen, J.A. (1999) The influence of the SW monsoon upon the nutrient biogeochemistry of the Arabian Sea. Deep-Sea Research II, 46, 571–591. Wyrtki, K. (1971) Oceanographic Atlas of the International Indian Ocean Expedition, National Science Foundation, Washington, DC. Yoshinari, T., Altabet, M.A., Naqvi, S.W.A., Codispoti, L.A., Jayakumar, A., Kuhland, M. and Devol, A.H. (1997) Nitrogen and oxygen isotopic composition of N2O from suboxic waters of the eastern tropical North Pacific and the Arabian Sea – measurement by continuous-flow isotope ratio monitoring. Marine Chemistry, 56, 253–264. Zietzschel, B. (ed.) (1973) The Biology of the Indian Ocean, Springer-Verlag, Berlin.

7

The northeastern Pacific abyssal plain Angelos K. Hannides and Craig R. Smith

7.1

Introduction

Abyssal plains occupy approximately two-thirds of the global sea floor, or two-fifths of the total surface of the earth. Owing to their large distance from the heavily populated continental margins, they are influenced only indirectly by land-based human activities, e.g. through atmospheric perturbations. However, planned in situ activities, such as the mining of manganese nodules, iron fertilization and commercial fishing may dramatically affect abyssal ecosystems in the future. In this chapter, we review the major geochemical and ecological characteristics of the northeastern Pacific abyssal plains, a comparatively wellstudied region by deep sea standards, and examine the system’s resilience in view of climatic change and in situ human activities. For the purposes of this review, we define the northeastern Pacific abyssal plain (NEPAP) as the oceanic region north of the equator and east of 180° W with waters deeper than 4000 m and slope angles less than 0.001 (Seibold & Berger, 1996). Although some recent authors consider the abyssal region in the world oceans to extend, ecologically speaking, from approximately 2000 to 6000 m depth (Gage & Tyler, 1991), the continental slope and the rise in the northeast Pacific (which reach ~4000 m) are ecologically distinct from the general abyssal plain. This is because the narrowness of the continental shelf and slope in the northeast Pacific yields steeper inclines and more energetic hydrodynamic conditions between 2000 and 4000 m depths than on the abyssal plains, allowing significant downslope transport of material from the continental margin (e.g. Smith & Demopoulos, 2003). Figure 7.1 and Table 7.1 indicate the region under consideration and major study sites and stations referred to here. 7.2

Key habitat parameters of deep seafloor communities

Deep seafloor communities are shaped by a number of key parameters that directly affect the nature and abundance of living organisms and their interactions with seafloor geochemistry. These parameters include (a) substratum type, (b) near-bottom hydrodynamic regime, (c) bottom-water oxygen concentration, (d) sinking particulate-organic-carbon (POC) flux, and (e) sediment redox conditions. Below, we describe these parameters and their variation in the northeastern abyssal Pacific.

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Fig. 7.1 Major study sites and stations in the northeastern abyssal Pacific are referred in this chapter (see Table 7.1 for details). The 4000 m bathymetric contour is also shown.

7.2.1

Key habitat parameters

7.2.1.1 Substratum type The type of substratum often determines, or at least is correlated with, trophic mode of many of the dominant animals, particularly the macro- and megafauna. Hard substrata are mostly inhabited by suspension feeders, whereas sediments are typically dominated by deposit feeders feeding primarily on surface sediments. Most of the deep seafloor is covered by sediments, consisting of silica and calcium carbonate tests of phytoplankton, and of terrigenous clays and other

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Table 7.1 Major study sites and stations in the northeastern Pacific referred to in this chapter Region

Site or study name

Site location

References

Eutrophic

EqPac transect

5° S–5° N 140° W

PACFLUX II FLUPAC TSII MANOP station C

0–1° N 103–140° W 0° N 150° W 1° N 139° W

Honjo et al. (1995); Hammond et al. (1996); Berelson et al. (1997); Smith et al. (1997); Brown et al. (2001) Hammond et al. (1996) Dunne et al. (2000) Cochran (1985); Dymond and Collier (1988)

EqPac transect

5–9° N 140° W

ENP DOMES A

5° N 125° W 8°27′ N 150° 47′ W

MANOP station S

11° N 140° W

PRA

12°57′ N 128° 19.5′ W

ECHO

14°40′ N 125° 26′ W

NIXO 47

14°0–40′ N 130° 40–60′ W

ALOHA-HOT

22°45′ N 158° W

Climax II CNP/MPG-I

Diameter of 50 km centered on 28°28′ N 155°20′ W 30–32° N 157–159° W

G

34° N 133° W

Mesotrophic

Oligotrophic

Honjo et al. (1995); Hammond et al. (1996); Berelson et al. (1997); Smith et al. (1997); Brown et al. (2001) Mullineaux (1987) Paterson et al. (1998); Glover et al. (2002) Cochran (1985); Dymond and Collier (1988) Paterson et al. (1998); Glover et al. (2002) Paterson et al. (1998); Glover et al. (2002) Renaud-Mornant and Goubault (1990) Glover et al. (2002); Smith et al. (2002) Hessler and Jumars (1974); Paterson et al. (1998) Smith et al. (1983); Snider et al. (1984); Mullineaux (1987); Smith (1987); Smith (1992) Smith et al. (1983); Smith (1987)

mineral grains transported by turbidites or by winds (Seibold & Berger, 1996). Hard substrata are limited to regions of recent volcanic activity, such as mid-ocean ridges, submarine canyons, seamounts and volcanic islands, and areas of very low sedimentation conditions, which allow the accretion of ferromanganese nodules (Ghosh & Mukhopadhyay, 2000; McMurtry, 2001). 7.2.1.2 Near-bottom currents Current regimes are also correlated with feeding mode. Very slow currents yield low horizontal particle flux and allow even fine particles to settle. This, combined with the absence of primary production in the dark abyss, yield very low concentrations of suspended food particles, often making suspension

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feeding unsustainable. High currents can cause sediment erosion and transport, disrupting most feeding activities and burying slow-moving organisms (Aller, 1989). Intermediate currents allow for a diversity of feeding modes and redistribute organic matter (e.g. phytodetritus) deposited on the sediment surface, causing spatial and temporal heterogeneity in resources (Lampitt, 1985; Smith et al., 1996; Bett et al., 2001). 7.2.1.3 Bottom-water oxygen Oxygen is a metabolic requirement for all metazoans and for aerobic metabolism in microbial taxa. Below a concentration of approximately 0.5 ml l−1, oxygen availability may affect community structure by excluding less tolerant taxa (Diaz & Rosenberg, 1995; Levin & Gage, 1998). Deep sea regions underlying high productivity areas and/or bathed in old water masses (i.e. water masses not recently in contact with the atmosphere) are particularly likely to experience low oxygen concentrations (Diaz & Rosenberg, 1995). 7.2.1.4 Sinking POC flux Organic matter synthesized at the ocean’s surface through photosynthesis constitutes the primary source of energy for seafloor communities (excluding hydrothermal vents and cold seeps). During sinking, most of the photosynthesized POC decomposes in the water column (Field et al., 1998; Hartnett et al., 1998). The amount of decomposition depends, in part, on oxygen presence, sinking velocity and water column depth (Suess, 1980; Jahnke, 1990; Hartnett et al., 1998). Sinking matter containing ballast minerals such as silica and calcium carbonate will sink faster and therefore may dominate POC flux to the deep ocean (Armstrong et al., 2001). Because very little of the primary production in the surface ocean reaches the abyssal seafloor (typically a few percent), the seafloor communities in abyssal regions are among the most food- and biomass-poor on the planet. The total biomass of many components of abyssal benthic communities appear to be correlated with annual POC flux (Rowe et al., 1991; Smith et al., 1997), whereas only microbes have been shown to respond to seasonal fluctuations in POC flux (reviewed by Gooday, 2002). 7.2.1.5 Redox conditions The most common reduction reactions taking place in the deep sea are aerobic respiration, denitrification, and iron and manganese reduction. In this respect, deep sea sediments differ from most coastal sediments. This difference is attributed to relatively high POC flux and sediment organic-carbon content in shallow water, where organic-carbon decomposition depletes oxygen and, in turn, other thermodynamically preferred microbial electron acceptors (in sequence, nitrate, iron, manganese), leaving sulphate reduction as the major metabolic process (Berner, 1980). In abyssal sediments, labile POC typically

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becomes depleted prior to the advent of sulphate reduction. Nonetheless, along gradients of POC flux to the abyssal seafloor, the relative importance of the various electron acceptors in microbial metabolism, and hence sediment redox conditions, will vary (e.g. Hammond et al., 1996). 7.2.2

Variation of key habitat parameters in the northeastern Pacific abyssal plain

7.2.2.1 Sediment types Most of the NEPAP is covered with soft sediments. The character of these soft sediments is largely determined by the surface productivity of the overlying ocean. Within the equatorial divergence zone, sediments shallower than ~4600 m are rich in calcium carbonate (50–90% by weight) and poor in organic carbon (<0.3% by weight) because of the large contribution of foraminiferan and pteropod tests (Jahnke, 1996); at greater depths, sediments are primarily siliceous muds comprised of diatom and radiolarian tests. Beyond the equatorial zone of high productivity (spanning roughly 5° S–5° N at 140° W), aeolian red clay particles (typically <6 μm in diameter) of terrigenous and volcanic origin dominate abyssal sediments (Hessler & Jumars, 1974; Smith et al., 1983). These sediments are poor in organic matter (<0.25% organic carbon by weight; Hessler & Jumars, 1974) relative to siliceous muds, which contain 0.25–0.5% of organic carbon by weight (Berger, 1974). The dominant hard substratum in the NEPAP is the exposed surface of ferromanganese nodules. These nodules precipitate over millions of years in areas of very low sedimentation, and range in size from 0.5 to 20 cm (McMurtry, 2001). In regions of high occurrence, such as between the Clarion and Clipperton fracture zones, they may cover more than 75% of the seafloor, whereas in sediments underlying the North Pacific oligotrophic gyre, they cover approximately 30% (Hessler & Jumars, 1974; Mullineaux, 1987; Ghosh & Mukhopadhyay, 2000). 7.2.2.2 Near-bottom currents and oxygen concentrations Most of the deep waters in contact with the abyssal sediments are a mixture of North Atlantic and Antarctic bottom-water masses. They are relatively saline and fairly cold (0.5–1.5°C). Currents at these depths typically are slow and do not impose shear stress of sufficient magnitude to transport sediments (Gardner et al., 1984; Demidova, 1999). Although bottom-water masses are relatively old in the NEPAP, bottomwater oxygen concentrations remain well above 2 ml l−1 and hence they probably do not influence benthic community structure (Levin & Gage, 1998). In the eastern tropical Pacific, oxygen-minimum zones appear in mid-water (100–1000m) due to intense water column decomposition but they do not extend to abyssal depths (Wishner et al., 1990).

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7.2.2.3 POC flux and redox conditions Due to the extreme food limitation of the NEPAP and its isolation from many other potential sources of energy and matter, POC flux constitutes a critical variable determining a multitude of habitat characteristics, such as sediment type (oozes and the presence or absence of ferromanganese nodules), redox conditions, rates of biogeochemical processes, and the biomass and structure of benthic communities (Hammond et al., 1996; Levin & Gage, 1998; McMurtry, 2001). Two major conditions determining POC flux to aphotic marine sediments are surface primary production and sinking time (Field et al., 1998). Since the depth of the water column within the NEPAP is relatively constant, any observed variations in most of the above characteristics will be determined by variations in surface primary productivity, or more accurately, export production, at each location. Based on POC flux to the sediments, then, the NEPAP roughly along 140° W can be divided into three zones: (a)

(b) (c)

the eutrophic abyss, from the equator up to 5° N, where the POC flux is ~1–2 g C m−2 y−1 (Berelson et al., 1997; Dymond & Collier, 1988; Honjo et al., 1995). This zone is restricted to 1–2° N as one moves west towards 180° W (Jahnke & Jackson, 1992), the mesotrophic abyss, from 5° N to 15° N, where the POC flux is ~0.5–1.6 g C m−2 y−1 (Dymond & Collier, 1988; Honjo et al., 1995), and the oligotrophic abyss, underlying the central North Pacific gyre, where POC fluxes typically are lower than 0.5 g C m−2 y−1 (Smith, 1987; Smith et al., 2002).

In consequence, seafloor geochemical conditions across these different zones co-vary with primary production rates and POC fluxes to the sediments. Profiles of dissolved O2 in sedimentary porewater exhibit deeper oxygen penetration with increasing distance from the equator (Hammond et al., 1996). This pattern is explained by the observed oxygen consumption by metabolism within sediments, which decreases from 0.24–0.6 mol m−2 y−1 in the eutrophic abyss to less than 0.06 mol m−2 y−1 in the oligotrophic abyss (Hammond et al., 1996; Jahnke & Jackson, 1992). As a consequence, oxygen becomes significantly depleted within 4–6 cm of the sediment–water interface within eutrophic abyssal sediments, suggesting an increasing importance of anaerobic decomposition in eutrophic abyssal sediments relative to oligotrophic sediments (Hammond et al., 1996).

7.3

Northeastern Pacific abyssal zones

Due to the overarching significance of POC flux in shaping abyssal communities, our discussion of the Northeastern Pacific abyssal plain ecosystem is organized into sections on the eutrophic, mesotrophic and oligotrophic abyss.

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The gradients in POC flux across the NEPAP explain gradients in the total benthic biomass, sediment community oxygen consumption, other solute fluxes into and out of the sediments and, consequently, elemental cycling rates. 7.3.1

The eutrophic equatorial abyss

The eutrophic equatorial abyss is characterized by high primary productivity of surface waters within the equatorial divergence zone and POC fluxes to the abyssal seafloor of ~1–2 g C m−2 y−1 (Dymond & Collier, 1988; Honjo et al., 1995; Berelson et al., 1997). The region of the abyssal seafloor impacted by this productivity extends to 5° N and S at 140° W, gradually narrowing as one moves west. At the 180th meridian, the eutrophic abyss extends only to 1–2° N and S (Jahnke, 1996; Jahnke & Jackson, 1992). This yields an area of approximately 4.5 million square kilometers that annually receives approximately 8000 Gt of POC from surface waters. At times, about 3% of this flux may be recently deposited phytodetritus (Smith et al., 1996), which may constitute an important resource for surface-deposit-feeding epifauna not observed in other zones. The composition of the living community is summarized in Table 7.2. The megafaunal community, and especially the burrowing fraction, remains poorly sampled. The dominant metazoan megafauna are urchins, hexactinellid sponges of the genus Hyalonema and epibenthic holothurians (Hoover, 1995; Smith & Demopoulos, 2003). Various protozoan agglutinating xenophyophores, especially of the genera Reticulammina and Stannophyllum, also are numerous and actually dominate megafaunal abundance (Smith & Demopoulos, 2003) but because of their high water content (~98%), they are unlikely to contribute greatly to biomass (Levin & Gooday, 1992). An indirect indicator of infaunal activities is feeding traces on the sediment surface and they suggest a significant presence of large echiurans (Smith & Demopoulos, 2003). Macrofaunal abundance and biomass are both dominated by polychaetes (62% and 62%, respectively), the rest being accounted for by tanaids, isopods and bivalves (Smith & Miller, in preparation). The vast majority (>90%) of individuals can be found at the top 3–5 cm of the sediment column (Smith & Demopoulos, 2003), an apparent response to localization of food resources near the sediment–water interface. Meiofauna are mostly represented by nematodes, which attain densities of 36–108 000 individuals m−2 and biomasses of 1.3–12 mg m−2 at the top 1 cm of sediment (Brown et al., 2001). Microbes, however, appear to contribute more to benthic biomass than macrofauna and meiofauna together. Where it has been quantified, microbial biomass may attain values of 0.2–0.3 g C m−2 (Smith et al., 1997), roughly equivalent to a biomass of 1.3–1.6 g m−2, assuming carbon constitutes approximately 16% of total biomass (Fenchel et al., 1998; Madigan et al., 2000). Most of the polychaetes and nematodes in the eutrophic abyssal zone belong to families considered to be deposit feeders (e.g. Kukert & Smith,

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Table 7.2 Benthic community components in the NEPAP. Large ranges in reported values may occasionally be due to differences in sediment depths analyzed. When possible, these differences have been normalized to a certain depth across regions Ecosystem type Habitat parameter

Eutrophic abyss

Mesotrophic abyss

Oligotrophic abyss

Sedimentary POC flux (g C m−2 y−1) Megafaunal abundance (ind. m−2) Megafaunal biomass (g wet wt m−2) Macrofaunal abundance (ind. m−2) Macrofaunal biomass (mg wet wt m−2) Meiofaunal abundance (103 ind. m−2) Meiofaunal biomass (mg wet wt m−2) Microbial abundance (1012 ind. m−2) Microbial biomass (mg wet wt m−2) Manganese nodule faunal abundance (103 ind. 0.25 m−2)

1.53–1.97(1)–(3)

0.26–1.65(2),(3)

0.04–0.76(4),(5)

0.17–0.25(6),a

0.10–0.17(6),b

0.15(7)

Not available

Not available

>12.4–12.6(7),c

1200–2000(6),(8),(9)

60–1200(6),(8),(10)

12–160(5),(10),(11)

400–600(6),(9)

120–400(6)

2.1–137(5),(7)

>36–108(12),d

>23–189(12),(13),d

10–232(11),(12),(14),e

>2.26–20.20(12),f

>2.07–17.01(13),g

0.24–243(11),(12),(14)

20–25h

13–20h

0.56–2.4(4),(14)

1250–1563i

813–1250i

95–172(4),(14)

–

1.05–1.40(15)

0.7–1.0(15)

With xenophyophores, 1.9–5.9 ind. m−2 (Smith & Demopoulos, 2003). With xenophyophores, 2.35 ind. m−2 (Smith & Demopoulos, 2003). c Estimated ash-free dry weight (AFDW) of the dominant holothurians and sea anemones at Station CNP back-calculated from reported AFDW carbon mass per unit area (Smith, 1992). d Data available only for nematodes at 0–1 cm. e Data available for foraminifera and metazoa, including nematodes. f Biomass data available only for nematodes at 0–1 cm; projected to 0–5 cm using reported contribution of 0–1 cm meiofaunal biomass to 0–5 cm meiofaunal biomass: average of 59.5% for 0–5° N and 50% at 9° N (Brown et al., 2001). g Meiofaunal biomass estimated using reported biomasses at Station NIXO 47 (Renaud-Mornant & Goubault, 1990) and corrected for an oligotrophic deep sea site (wet wt/nematode = 0.09 μg; Vanreusel et al., 1995). h Calculated using reported microbial biomass of 0.2–0.25 g C m−2 for 0–5° N and 0.13–0.2 g C m−2 for 5–9° N (Smith et al., 1997), and a conversion factor of 10 fg C/cell (10 × 10−15 g C/cell) (Karl & Dobbs, 1998). i Calculated using reported microbial biomass of 0.2–0.25 g C m− 2 for 0–5° N and 0.13–0.2 g C m−2 for 5–9° N (Smith et al., 1997), and a % wet weight carbon content of microbes of 16% (Fenchel et al., 1998; Madigan et al., 2000). References: (1) Berelson et al. (1997); (2) Dymond and Collier (1988); (3) Honjo et al. (1995); (4) Smith et al. (2002); (5) Smith (1987); (6) Smith et al. (1997); (7) Smith (1992); (8) Glover et al. (2002); (9) Smith and Miller (in preparation); (10) Paterson et al. (1998); (11) Hessler and Jumars (1974); (12) Snider et al. (1984); (13) Brown et al. (2001); (14) Renaud-Mornant and Goubault (1990); and (15) Mullineaux (1987). a

b

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1992). Their activities contribute to the relatively high abyssal bioturbation rates in this region. Biodiffusion coefficients at stations of the JGOFS EqPac transect determined using 234Th and 210Pb are correlated with POC flux, and decrease by an order of magnitude from the eutrophic to the mesotrophic abyss (Smith et al., 1997). The depth of mixing in the eutrophic zone extends down to 6–8 cm (Smith etal., 1997; Smith & Rabouille, 2002). Biodiffusion coefficients estimated using the short-lived 234Th are much higher than those estimated using 210Pb, which suggests age-dependent mixing, that is the preferential bioturbation of newly deposited particles (Smith et al., 1993). Additionally, the activity of burrowing urchins appears to homogenize surface sediments down to 3 cm, a feature recognizable in 210Pb profiles (Hoover, 1995; Smith et al., 1997). Diagenetic modeling has aided in the estimation of organic matter residence times in eutrophic abyssal sediments. Two fractions appear to be present: a labile fraction with a residence time of 7–150 days, and a refractory fraction with residence time between 44 and 289 years (Hammond et al., 1996). The average time needed for surface-deposited organic matter to reach a certain depth in the sediment may be calculated using the average depth of the mixed layer and the average bioturbation coefficient. At the eutrophic NEPAP, using a bioturbation coefficient of 0.31 cm2 y−1 from 210Pb profiles (0–5° N; Cochran, 1985; Smith et al., 1997), it is estimated that reaching 1 cm into the sediments takes 3.2 years, while reaching the bottom of the mixed layer at 6–8cm (Smith & Rabouille, 2002) may take 116–206 years. This suggests that benthic organisms living on or within 1cm from the sediment surface may show very little resistance and resilience to changes in POC flux, relative to the infauna living deeper in the sediment. Benthic oxygen flux into the equatorial abyssal sediments varies between 0.24 and 0.6 mol m−2 y−1 and is correlated roughly to the supply of organic matter from above, decreasing as one moves away from the equator (Jahnke & Jackson, 1992; Hammond et al., 1996). Roughly 90% of this oxygen flux results from the decomposition of the labile fraction of organic matter. As inferred from flux data, eutrophic abyssal sediments are the site of regeneration and release of various nutrients including nitrate, phosphate and silica (Hammond et al., 1996). Finally, the high oxygen consumption rates (Jahnke & Jackson, 1992) and steep porewater oxygen profiles (Hammond et al., 1996) in the equatorial abyss suggest that anaerobic processes may occasionally constitute significant metabolic pathways compared to meso- and oligotrophic sediments. In general, this region demonstrates high organic-matter processing rates and rapid elemental cycling by abyssal standards, performed to a large extent by abundant microbiota. The flux of organic matter from the surface appears to be high enough to support a significant surface-deposit feeding megafaunal community, and densities and biomasses of macrofauna and meiofauna that are high by abyssal deep sea standards. Diversity has only been determined for certain taxa and shown to be high on local scales. For example, within the

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polychaetes, there may be up to 83 species per 163 individuals (Glover et al., 2002). However, the very small sampled area does not allow reliable estimates of diversity on regional scales. Because patterns of biodiversity and their causes are so poorly constrained in the eutrophic abyss, it is not fruitful to speculate on biodiversity responses to ecosystem change. 7.3.2

The mesotrophic (sub-equatorial) abyss

The mesotrophic (or sub-equatorial) abyss constitutes a zone of transition between the region of high POC beneath equatorial upwelling and the regions of very low fluxes beneath the oligotrophic central gyres. North of the equator roughly along 140° W, this zone stretches between approximately 5° and 15° N, or for 10° N of the boundary of the eutrophic abyss. A major characteristic of this region is the concentration of manganese nodules dotting vast areas of the seafloor, covering up to 75% of its plan area (Ghosh & Mukhopadhyay, 2000). The surrounding sediments contain between 0.25 and 0.5% organic carbon and significantly less CaCO3 than the equatorial abyssal sediments (Berger, 1974; Jahnke, 1996; Smith et al., 1997). The mesotrophic abyssal benthic standing stocks are summarized in Table 7.2. One major difference between the eutrophic and mesotrophic abyss is a large reduction in the abundance of surface-burrowing urchins and echiuran worms, indicated by the disappearance of urchin furrows and spoke traces (Hoover, 1995; Smith, unpublished data). Protozoan xenophyophores do remain common (approximately 2.35 m−2; Smith & Demopoulos, 2003), but metazoan megafaunal abundance drops to almost half that of the eutrophic abyss (Table 7.2; Smith etal., 1997). The metazoan megafauna appear to be dominated primarily by hexactinellid sponges of the genus Hyalonema (55–87%), and secondarily by surface-deposit feeding holothurians (Hoover et al., 1994; Smith & Demopoulos, 2003). Macrofaunal abundance and biomass are diminished even further, down to a third of those in the equatorial abyss (Table 7.2). A similar pattern is observed for meiofaunal biomass and abundance (Brown et al., 2001), which in this region are dominated by nematodes followed by copepods and foraminifera (Renaud-Mornant & Goubault, 1990). Microbial biomass also declines (Table 7.2); however, the decline in microbial biomass is smaller than in the larger size classes, yielding a greater contribution of microbes to total mesotrophic community biomass (Smith et al., 1997). The vast majority of macrofauna in the mesotrophic abyss have been determined to be deposit feeders (Paterson et al., 1998), whereas meiofaunal abundances are correlated with microbial biomass, suggesting that they may be bacterial grazers (Brown et al., 2001). Manganese nodules in this region (site ENP) harbor sessile eukaryotic epibionts, with 98.2% of the abundance attributed to foraminifera and rhizopod

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protozoans, with sponges, mollusks, polychaetes, bryozoans and nematodes also present (Mullineaux, 1987). Although nodule-associated nematodes are not as abundant as in the surrounding sediments, they appear to be diverse (Mullineaux, 1987). Accompanying the decrease in POC flux relative to the eutrophic abyss is a decrease in sedimentary oxygen consumption rates varying between 0.06 and 0.24 mol m−2 y−1 (Jahnke & Jackson, 1992; Hammond et al., 1996). Nutrients such as nitrate, phosphate and silica are regenerated from sediments at rates lower than those of the eutrophic abyss, as predicted by differences in POC supply (Hammond et al., 1996). The presence of a labile, short-lived organic matter fraction predicted by modeling in eutrophic sediments has not been verified for mesotrophic sediments. Instead, only one relatively refractory pool of organic matter appears to be present with estimated residence time of 9–50 years (Hammond et al., 1996). As in the previous section, using an average bioturbation coefficient of 0.19 cm2 y−1 (5° N; Smith et al., 1997) and an average mixed layer depth of 4 cm (Smith & Rabouille, 2002) for the mesotrophic Pacific, it is estimated that it may take an average of 84 years for freshly deposited organic matter to reach the bottom of the mixed layer. The shallower mixed layer and the consequent shorter time period needed for mixing suggest that benthos at the mesotrophic abyss may be even less resilient than that of the eutrophic abyss to changes in POC fluxes. On local scales, species diversity in the mesotrophic abyss is lower than in the eutrophic abyss. On average, 45–75 different species of polychaetes have been identified for every 150–163 individuals examined (Paterson et al., 1998; Glover et al., 2002), and 780 nematode individuals could belong to as many as 45 species (Renaud-Mornant & Goubault, 1990). Because of the high abundance of nickel- and cobalt-rich manganese nodules in this region, mining of nodules from this area has been contemplated since the 1950s (Glasby, 2000). The paucity of ecological and biogeographic information from this region, e.g. concerning colonization rates of benthos and species ranges, constitutes a major hurdle in assessing the potential environmental impacts of nodule mining. 7.3.3

The oligotrophic central gyre abyss

A large proportion of the NEPAP features some of the lowest POC fluxes and lowest biological standing stocks and rates recorded at the seafloor. POC fluxes in the northeastern Pacific oligotrophic abyss are typically lower than 0.5 g C m−2 y−1 (Smith, 1992; Smith et al., 2002), although some seasonality is present as demonstrated by massive pulses of organic matter of up to 2 mg C m−2 d−1 to the seafloor (Karl, unpublished data; Karl et al., 1996; Smith et al., 2002). The very low supply of organic matter to the seafloor results in deep penetration of oxygen into sediment porewaters (Hammond et al., 1996) and very low sedimentary oxygen consumption rates (<0.06 mmol m−2 y−1;

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Jahnke & Jackson, 1992). Sediments reflect these conditions, consisting mostly of red clay particles smaller than 6 μm transported on winds from the continents (Hessler & Jumars, 1974; Smith et al., 1983), and carrying very little organic matter, typically less than 0.25% by weight (Berger, 1974; Smith et al., 1983). The oligotrophic conditions yield very low biomass and abundance for all components of the benthic community (Table 7.2). The megafauna consists of epibenthic holothurians (primarily Amperima sp.), cnidarians and xenophyophores at densities about half those of the eutrophic abyss. In spite of the oligotrophic conditions, bottom-associated mobile scavengers such as lysianassid amphipods, rattail fish and decapods appear to be important components of this community (Dayton & Hessler, 1972; Ingram & Hessler, 1983; Smith et al., 1992; Priede et al., 1994). These scavengers are not typically encountered during surveys of the seafloor, but are attracted to bait-fall experiments within minutes to hours and form large aggregations (tens to hundreds per fall) that consume 10–100 kg of carrion within days (Dayton & Hessler, 1972; Ingram & Hessler, 1983; Priede et al., 1994). Macrofaunal assemblages are dominated by polychaetes (55% of individuals) with other major components being tanaids (18%), bivalves (7%) and isopods (6%) (Hessler & Jumars, 1974). Macrofaunal biomass and abundance are substantially lower than those of the eutrophic abyss, and the average macrofaunal body size is approximately an order of magnitude smaller than that of the eutrophic and mesotrophic abyssal regions (Smith & Demopoulos, 2003). Meiobenthic biomass at the CNP/MPG-I site is dominated by foraminifera (87%), which account for 50% of the total abundance, with nematodes contributing 45% (Snider et al., 1984). The microbial community dominates biomass of total benthos in the oligotrophic abyss, exceeding that of macrofauna and meiofauna combined by approximately ten times (Table 7.2). Moreover, microbial biomass appears to fluctuate seasonally with POC flux, unlike other community groups (Smith et al., 2002). Manganese nodules in this region (station CNP) host assemblages similar to the mesotrophic region (station ENP), with foraminifera and rhizopod protozoans accounting for 99.5% of the abundance (Mullineaux, 1987). Considering the scarcity and high degree of specialization of scavengers in the oligotrophic abyss, the major source of energy for the abyssal communities appears to be sinking particulate matter, and due to the very low fluxes, the oligotrophic abyss appears to be food-deficient (Smith, 1992; Smith et al., 2002). The majority of macrobenthos and meiobenthos are deposit feeders (Hessler & Jumars, 1974), while the very low community metabolic rates are presumably dominated by the microbial component (Smith, 1992). Low standing stocks and relatively high species diversities (more than 45 species for every hundred polychaete individuals; Hessler & Jumars, 1974; Glover et al., 2002) suggest that the study of metazoan life histories, which requires collection of

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large numbers of individuals per species, will be particularly difficult in the oligotrophic abyss. Knowledge of these life histories is essential in addressing impacts of global change and anthropogenic disturbance in all abyssal regions. 7.4 7.4.1

Sensitivity and resilience to natural and anthropogenic change General thoughts

The NEPAP ecosystems are characterized by extremely low energy inputs and elemental cycling rates, very low biological process rates, high species diversity and, usually, high physical stability. The structure of the community and the intensity of cycling rates at any one abyssal location are predominantly determined by the POC flux regime at that location. As a consequence, these abyssal ecosystems are, probably, extremely sensitive to long-term changes in POC fluxes; even modest sustained changes may yield substantial alterations in community structure. For example, if eutrophic to oligotrophic trends are representative, a three-fold drop in mean POC export to the deep sea may yield a 10-fold decline in macrofaunal and microbial abundance and biomass (Table 7.2). However, these abyssal ecosystems may be surprisingly resilient (i.e. change little) when faced with short-term variations in POC flux. This is because POC flux to the NEPAP seafloor can vary substantially within and between years, e.g. due to ENSO events and phytodetrital pulses (Dymond & Collier, 1988; Honjo et al., 1995; Smith et al., 1996; Smith et al., 2002). In addition, abyssal species are likely to (1) be well adapted to extended periods of low food availability, and (2) have very low growth and reproduction rates (Gage & Tyler, 1991). As a consequence, flux-induced changes in macrofaunal community structure may occur relatively slowly, i.e. over at least several years. Thus, we expect that the NEPAP ecosystems would ultimately exhibit high sensitivity to moderate, but long-term, changes in POC flux; however, the consequent changes in benthic community structure would probably occur slowly. In contrast, benthic community structure and processes in the NEPAP region are likely to be both extremely sensitive to, and have very little resistance to, physical perturbations (e.g. mining disturbance). This is because the natural ecosystem is relatively very stable (compared to virtually all other ecosystems), most animals are small and/or very delicately constructed, and critical habitat structure for the entire benthic fauna is concentrated within a few centimeters of the sediment–water interface. Thus, it would require very little physical energy to disrupt the animals and the thin veneer of surface sediments that define this ecosystem. The extremely low sediment accumulation rates, bioturbation rates, nodule growth rates and macrofaunal recolonization rates of the NEPAP seafloor ecosystem, compared to other seafloor habitats (Smith & Demopoulos, 2003), suggest that recovery from physical disturbance is likely to be extremely slow relative to other ecosystems.

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Global biogeochemical cycles may respond to natural or anthropogenic changes such as variations in El Niño–Southern Oscillation (ENSO) frequency or global warming, and in the process cause perturbations in oceanic ecosystem functions including primary production and the export of POC to the abyssal ocean. Human activities on the abyssal seafloor and in the overlying water column, e.g. mining of ferromanganese nodules and iron fertilization, could impact abyssal communities more rapidly and dramatically than atmospheremediated changes. In the following paragraphs, we discuss potential impacts of a number of natural and anthropogenic influences on NEPAP benthic communities. 7.4.2

Potential sensitivity and resilience to specific changes

7.4.2.1 Climate variation in the equatorial and North Pacific A number of large-scale climatic cycles and their impacts have been recognized and thoroughly studied in the 1990s. Two of the climatic cycles that impact the central and northern Pacific are the ENSO and the Pacific Decadal Oscillation (PDO). ENSO events are characterized by reversals of the atmospheric pressure systems in the south and central Pacific and switching of the sources of the intensely upwelled water masses along the southeastern Pacific continental margin. An El Niño event results in warmer surface waters, greater stratification and reduced upwelling of nutrients in the central and eastern equatorial Pacific, reducing primary productivity. While reduced equatorial–subequatorial productivity during El Niño is observed in the central eastern Pacific in all studied events (Barber & Chavez, 1983; Barber et al., 1996; Strutton & Chavez, 2000), the effects of ENSO events on POC flux to the seafloor are not clear. At least three studies from the central equatorial Pacific contrast deep POC fluxes under El Niño and non-El Niño conditions. Dymond and Collier (1988) explored the effects of the 1982 El Niño on POC export at MANOP stations C and S (Fig. 7.1 and Table 7.1). They found decreased export to the eutrophic abyss during the El Niño event, but saw an increase in export to the mesotrophic abyss at 11° N over the same time period (Fig. 7.2). POC flux data were also recorded along the US JGOFS EqPac transect during and after the moderate El Niño event of 1991–92 (Honjo et al., 1995; see Murray et al., 1995 for study description). The POC fluxes at 700 m above the bottom during El Niño were higher at the equator, 2° N and 9° N but lower at 5° N, 2° S and 5° S relative to non-El Niño conditions (Fig. 7.2). However, the French JGOFS FLUPAC and Zonal Flux studies reported no statistical differences between POC fluxes calibrated using 234Th out of the euphotic zone (150 m) at 0° N 150° W during the following El Niño event of 1994 (FLUPAC TSII) and during non-El Niño conditions (Zonal Flux) (Dunne et al., 2000). Considering the evident increase in frequency of El Niño events during the latter part of the

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2.5

Corg (g m–2 y–1)

B A

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1.5 A 1.0 B 0.5 0.0 0

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Fig. 7.2 Deep sea POC fluxes in the equatorial and sub-equatorial Pacific during El Niño and non-El Niño periods. All sediment trap stations are along 139–140° W. Data for latitudes 1° N (MANOP station C, 2908–3495 m) and 11° N (MANOP station S, 3400 m) are from Dymond and Collier (1988) for the 1982–83 event (marked A). Data for latitudes 0° (3618 m), 2° N (2200 m), 5° N (3800 m) and 9° N (4400 m) along the US JGOFS transect are from Honjo et al. (1995) for the 1991–92 event (marked B). Courtesy of R. Collier, US JGOFS, http://usjgofs.whoi.edu/jg/dir/jgofs/eqpac/.

20th century (Trenberth & Hoar, 1996, 1997; Timmermann et al., 1999) and the defining role of POC fluxes on NEPAP communities, the impact of ENSO events on the seafloor is yet to be resolved. The PDO climate pattern appears to be teleconnected to ENSO and is characterized by interdecadal climate shifts in the North Pacific, as observed using a variety of environmental indices from sea surface temperatures to fish productivity in the region (Mantua et al., 1997; Minobe, 1999). Two recent climate-regime shifts in the North Pacific in 1976–77 and in 1989, have been explored and documented extensively (Ebbesmeyer et al., 1991; Hare & Mantua, 2000). North Pacific climate indices generally correlate very well with both North Pacific environmental indices and ENSO indices (Mantua et al., 1997), and are expected to affect, among other things, primary productivity in the North Pacific sub-tropical gyre. Past productivity changes have been suggested to occur due to variation in mixed-layer depth over decadal scales (Polovina et al., 1994, 1995) and lengthened periods of stratification, and reduced turbulence (Karl et al., 1995, 2001a). In particular, it has been suggested that chlorophyll concentrations in the euphotic zone at stations Climax and ALOHA document the impact of PDO, and specifically the climate shift of 1976–77 on oligotrophic gyre productivity (Karl, 1999; Karl et al., 2001b). The current hypothesis suggests a shift from a eukaryotic phytoplankton community to one dominated by nitrogen-fixing prokaryotes. The shift to prokaryotic dominance is caused by physical conditions

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that restrict nutrient supply to the euphotic zone and encourage nitrogen fixation either by free-living cyanobacteria or by diatom endosymbionts (Karl et al., 1995, 2001a; Zehr et al., 2000). Due to the presence of these nitrogen-fixing endosymbionts, the otherwise plausible translation of this regime shift into a decrease in diatom and coccolithophorid populations and productivity, and consequently POC export changes (Smith et al., 2002), is not possible. Time series data from deep sea sediment traps at station ALOHA (D.M. Karl, unpublished data) have continuously been collected well after the last proposed regime shift in 1989 (Hare & Mantua, 2000), and they may prove useful in interpreting the role of community structure changes on POC export from the ocean surface layers on the occasion of the next shift. The effects of ENSO and PDO on communities of the NEPAP are far from being resolved. Whereas their impacts in the regions immediately affected (equatorial and sub-equatorial Pacific for ENSO and northeastern Pacific for PDO) can be speculated, their impacts on larger-scale regional conditions appear to be synergistic for some regions and opposing for others (e.g. Mestas-Nuñez & Enfield, 2001). For example, it has been suggested that subtropical waters of the Pacific experience increased productivities during El Niño events whereas equatorial productivity drops (Leonard et al., 2001). In any case, from the data available at present, it appears that all regions of the NEPAP are experiencing interannual to interdecadal variations in export due to large-scale climate patterns. During El Niño events, zones of the eutrophic abyss may be starved whereas zones of the mesotrophic abyss may be enriched. This pattern will probably be longer lasting, considering the increased frequency of ENSO. As a result of larger-scale climatic teleconnections, the oligotrophic abyss is undergoing similar starvation-enrichment oscillations on the frequency of decades. 7.4.2.2 Global increase in atmospheric greenhouse gases and temperatures Fossil fuel burning is now accepted to have increased the concentrations of greenhouse gases in the atmosphere and is expected to lead to global warming (IPCC, 2001). Global climate change may affect primary productivity directly, by physiological effects on the photosynthetic taxa dominating POC export and indirectly, by enhancement of regional climatic conditions which result in shifts in photosynthetic-community structure. In turn, alterations of primary production processes will impact the export of organic matter from the surface waters to the abyssal plain and the benthic biota that rely on this export. Mechanisms connecting changes in dissolved inorganic carbon (DIC) concentrations and water temperatures, and the physiology of photosynthesizing taxa dominating POC export, such as diatoms and coccolithophorids, are well studied. Marine diatoms have been shown to concentrate DIC using extracellular and intracellular carbonic anhydrase (Burkhardt et al., 2001; Tortell et al., 1997), thus avoiding the limitation of concentration-dependent diffusive uptake. Raven (1991) suggested that photosynthesizers which are now DIC-limited

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because they lack similar active uptake mechanisms may become more competitive if DIC concentrations increase, resulting in a change in the community structure of primary producers. Additionally, similar experiments performed with coccolithophorid species demonstrate a decrease in calcification rates and calcite:POC ratios with increasing DIC concentrations associated with a simultaneous increase in POC production rates (Riebesell et al., 2000). Data from station ALOHA suggest that the surface waters of the oligotrophic gyre have tracked the atmospheric increases in DIC for the last decade (Fig. 7.3; Karl et al., 2001b), although it is uncertain whether the magnitude of these changes is sufficient to alter the key primary producers or export production to the oligotrophic NEPAP. In contrast, surface water temperatures in the same region do not show any consistent pattern during the last decade and are presumably controlled by the

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Year Fig. 7.3 Trends in inorganic carbon concentrations between 1989 and 2001. Above: atmospheric carbon dioxide concentrations as measured at the Manoa Loa observatory. Data courtesy of the Carbon Dioxide Information Analysis Center, http://cdiac.ornl.gov/ndps/ndp001.html. Below: mixed layer (0–50 m; Karl & Lukas, 1996) DIC concentrations, normalized to a salinity of 35 per mil, at station ALOHA. Data courtesy of D.M. Karl, The Hawaii Ocean Time-series (HOT), http://hahana.soest.hawaii.edu (Karl et al., 2001b).

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climate cycles operating in the Pacific as described above. Nonetheless, higher water temperatures have been shown to decrease the size of most diatom species examined from the open ocean (Montagnes & Franklin, 2001). This matches widely observed patterns of decreasing ectotherm size with increasing temperature (see Atkinson, 1994 for a review). A reduction in size will result in a reduction in sinking rate, although the extent of the effect also depends on the associated changes in the cell’s specific gravity. If all cell components, including vacuoles and frustules, decrease proportionately then the specific gravity will not change, and any observed changes may be attributed only to size (Jackson, 1990; Montagnes & Franklin, 2001). Since the major seasonal POC fluxes at Station ALOHA are dominated mostly by diatoms during the late summer bloom (Scharek et al., 1999a,b) and by coccolithophorids during the winter– spring bloom (Cortés etal., 2001), they may be explained in part by inter-annual variations in average temperatures of the surface layers at this region (Fig. 7.4). Indeed, a fairly significant relationship between temperature and production export has been demonstrated to exist on a global scale and modeled (Laws et al., 2000) although it may not be robust at low primary productivity zones such as those overlying the NEPAP.

Average annual PC flux at z = 150 m (mg C m–2 d–1)

36 R2 = 0.4585 p = 0.022

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Average annual mixed layer temperature (°C) Fig. 7.4 Relationship between average annual particulate carbon flux at 150 m and mixed layer temperatures (0–50 m; Karl & Lukas, 1996) at station ALOHA. The relationship was tested using ANOVA (F = 7.62, p = 0.022). Data courtesy of D.M. Karl, The Hawaii Ocean Time-series (HOT), http://hahana.soest.hawaii.edu.

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Apart from physiological effects on major photosynthetic eukaryotes, global climate change has been connected to the increasing frequency of El Niño-La Niña cycles in the later part of the twentieth century (Trenberth & Hoar, 1996, 1997; Timmermann et al., 1999). Increased surface water stratification will certainly result in nutrient depletion, amplified importance of nitrogen fixation and variable changes in particulate matter export to the abyss as described above (Karl et al., 2001a). Modeling of the interaction between ENSO and global climate change suggests that not only the frequency but also the intensity of El Niño and La Niña events will increase in the future, if current trends in atmospheric carbon dioxide concentrations remain unchanged (Timmerman et al., 1999). Ultimately, the consequence of these changes may be a prolonged period of atrophic conditions in the abyssal plains subjecting the eutrophic and mesotrophic abyss to an energetic imbalance and a continuous demand for organic matter similar to that of the oligotrophic abyss (Karl, 1999; Smith et al., 2002). 7.4.2.3 Manganese nodule mining Manganese nodule mining may ultimately be the largest-scale human activity to directly impact the NEPAP or any other abyssal habitat worldwide (Thiel, 2001). Twelve pioneer investigator countries and consortia, including the International Seabed Authority, have carried out more than two hundred exploratory cruises around the globe to investigate locations of high manganese nodule coverage, especially the area between the Clipperton and the Clarion fracture zones (Fig. 7.5; Glasby, 2000). This region covers approximately 6 million km2 (out of a total of 46 million nodule-rich km2 worldwide) and is estimated to contain approximately 7.5 billion metric tons of manganese, 78 million tons of cobalt, 340 million tons of nickel and 265 million tons of copper (Ghosh & Mukhopadhyay, 2000; Morgan, 2000). At present, six contractors

km

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Fig. 7.5 Extent of the general Pacific area under prospecting for manganese nodule mining. The shaded sectors have either been allocated to, reserved for or claimed by pioneer investors and consortia. Maps courtesy of the International Seabed Authority, http://www.isa.org.jm.

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are licensed by the International Seabed Authority (the international body charged with managing deep sea mining) to explore nodule resources and to test mining techniques in six claim areas, each covering 75 000 km2 (i.e. an area roughly half the size of the state of Florida). When mining ultimately begins (probably not for another 10–20 years), each mining operation is projected to directly disrupt, through nodule harvesting, ~800 km2 of seafloor per year, and to disturb the sediment-dwelling fauna over an area 5–10 times that size due to redeposition of suspended sediments (data from the International Seabed Authority). Thus, in any given year, nodule mining might severely damage abyssal seafloor communities over areas of 20 000– 40 000 km2 (a zone of devastation at least the size of the state of Massachusetts). One obvious direct effect of manganese nodule mining will be removal of the nodules themselves, which will require millions of years to re-accrete (Ghosh & Mukhopadhyay, 2000; McMurtry, 2001). Nodule mining will thus remove the only hard substrate present at the NEPAP seafloor, yielding habitat loss and at least local extinction of the nodule fauna, which differs dramatically from that in surrounding sediments (Mullineaux, 1987). The nodule-collecting process will also undoubtedly remove at least the top 5 cm of sediment, broadcasting much of this material into the sediment column (e.g. Gage & Tyler, 1991). In the direct path of the collection device, most animals will be killed immediately and compacted relatively organic-poor subsurface sediments will be exposed (Jumars, 1981). Resuspended sediments will rain out onto the seafloor, burying benthos to varying depths, and covering the sediment–water interface with layers of subsurface sediments (Jumars, 1981). Because the NEPAP habitat is normally very stable, and is dominated by very small and/or fragile animals, the direct effects of nodule collection will be devastating for the benthos (Jumars, 1981). The indirect impacts of sediment redeposition may also be deleterious because much of the NEPAP macrofauna and megafauna appear to be surface-deposit feeders that consume the meager flux of POC from the water column. Redeposition of subsurface, presumably food-poor sediment from the mining plume may dilute food resources, causing nutritional stress to the benthos. This effect would be enhanced if redeposition were chronic, i.e., if it occurred over extended time periods (months to years) due to concentration of mining activities at a single site. A number of small-scale studies (relative to full-scale mining) have been conducted to evaluate the sensitivity and recovery times of abyssal benthic communities subjected to nodule-mining disturbance. These studies have used disturbances of the following two types: (1) A plowing disturbance, in which an 8 m wide frame with multiple plow heads was dragged over the ocean floor in the south equatorial abyss, directly disturbing about 20% of the seafloor within a circular area of 11 km2 [the DISCOL experiment (Thiel, 2001)] and (2) A sediment removal and redeposition disturbance, in which a sedimentpumping system has been towed along seafloor swathes 200–400 m wide and

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2–3 km long, removing sediments from a track ~10 cm deep and 2 m wide and disgorging them as at plume at an altitude of 5 m [the BIE-type experiments (Thiel, 2001)]. Both types of studies have addressed community disturbance and recolonization within directly disturbed tracks, as well as in areas subjected to rapid resedimentation to thicknesses of ≥7 mm. Although these experiments produced disturbances of much lower intensity and much smaller spatial scale than would result from commercial-scale mining, they provide some insights into the sensitivity and recovery times of abyssal nodule communities exposed to mining disturbance. Direct plowing disturbance (DISCOL) yielded a dramatic reduction in the abundances of macrofaunal polychaetes (51.4%), tanaids (72%), isopods (81.5%) and bivalves (90.7%) (Borowski & Thiel, 1998). Megafauna was also heavily impacted immediately after the disturbance, decreasing sharply in abundance within plow tracks (Bluhm et al., 1995). In areas of resettling sediment, the disturbances of macrofaunal and megafaunal abundance varied from quite low to negligible (Bluhm et al., 1995; Borowski & Thiel, 1998). Three years after the disturbance event, the abundance of dominant macrofaunal taxa, especially polychaetes, had reached pre-disturbance levels, but macrofaunal diversity remained depressed even after seven years (Borowski & Thiel, 1998; Borowski, 2001). The abundance and diversity of megafauna also remained below pre-disturbance levels seven years after disturbance (Bluhm, 2001). Nematode densities and biomass were not detectably affected by the disturbance event, while copepods only showed some decrease in diversity at the species level after seven years (Ahnert & Schriever, 2001; Vopel & Thiel, 2001). Results from sediment removal/redeposition (or BIE-type) experiments show similar sensitivity of abyssal seafloor communities to physical disturbance. Immediately following apparent redeposition of <1 cm of sediment (Yamazaki et al., 1997; Fukushima et al., 2000), reductions have been observed in the abundance of many components of the biological community, including megafauna (21–48%), macrofauna (up to 63%), meiofauna (23%) and microbiota (one to three orders of magnitude), possibly due to burial by settling sediment, and interference with respiratory and feeding functions (Fukushima et al., 2000; Ingole et al., 2001; Raghukumar et al., 2001; Rodrigues et al., 2001). Two years after redeposition, the abundance of surface-deposit-feeding and mobile megafauna may remain low, suggesting a lingering impact on their food resources (Fukushima et al., 2000). One must be very cautious in extrapolating these experimental results to the impacts of commercial mining, which would be much more intense and devastate much larger areas of seafloor (i.e. 100–1000 km2 versus 1–11 km2). One can only say that NEPAP benthos will be substantially disturbed by even modest amounts (~1 cm) of sediment redeposition resulting from mining activities, and that full sediment-community recovery from major mining disturbance will take much longer than seven years. We cannot predict the likelihood of

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species extinctions from nodule mining, because we do not know the typical geographic ranges of species living within the nodule region (are the ranges large or small relative to the potential spatial scales of mining disturbance?). We can predict that the recovery of the nodule fauna within a mined area will take millions of years because of the extremely slow regeneration rate of their required substrate, manganese nodules. It is clear that to fully predict commercial mining impacts, substantially more information is required concerning species ranges, burial sensitivity and the spatial-scale dependence of recolonization in the NEPAP biota. 7.4.2.4 Iron fertilization The goal of iron fertilization of the oceans would be to stimulate primary production and export production in iron-limited, high-nutrient-low-chlorophyll (HNLC) regions of the ocean, such as the eutrophic Pacific. This stimulation of the biological pumping of CO2 from the atmosphere to the deep ocean would be expected to reduce the atmospheric greenhouse effect, mitigating global warming (Fuhrman & Capone, 1991; Raven & Falkowski, 1999). The stimulation of primary production on a scale of 100–300 km2 with the addition of iron was demonstrated in the equatorial Pacific during the IronEx I and II experiments in 1993 and 1995, respectively (Martin et al., 1994; Coale et al., 1996) and during the SOIREE experiment in the Southern Ocean (Boyd et al., 2000). Where measured, the export of particulate carbon from the mixed layer did not increase significantly during these iron-enrichment experiments (days to weeks) (Coale et al., 1996; Trull & Armand, 2001). However, longer-term effects of iron fertilization on POC export were modeled for the SOIREE experiment (Hannon et al., 2001). The modeling suggested that POC export out of the fertilized patch of ~150 km2 may have increased by two- to threefold within two months after iron addition. Longer monitoring of POC export is needed in subsequent fertilization experiments to test these predictions. An extreme iron fertilization scenario modeled by Sarmiento and Orr (1991) relies on the sustained depletion of phosphorus in Southern Ocean surface waters by continuous iron addition. The three-dimensional, multi-layered model predicts a global increase of export production and the possibility of anoxia in deep oceanic waters. Specifically, the results indicate a POC export increase of 6–30 Gt C y−1, i.e., a doubling of export production, after 100 years of fertilization. In addition, anoxia is predicted for certain parts of the southwest Indian Ocean. Considering the enormous scale of this hypothesized fertilization operation (0.6 Mt utilizable Fe y−1; Sarmiento & Orr, 1991) and the low spatial resolution of the model, the exercise provides limited insights into more realistic iron fertilization scenarios. Because of the potential magnitude and uncertainty of environmental impacts, proposals to fertilize the ocean to mitigate global warming (or to earn carbon credits) have been controversial (e.g. ASLO, 2001; Chisholm et al., 2001;

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Johnson & Karl, 2002). Quite simply, existing models and experimental studies are inadequate to predict the ecosystem consequences of large-scale, prolonged iron fertilization. It is clear that if iron fertilization were successful on a large, enough scale to mitigate global warming, POC fluxes to large areas of the deep sea floor (possibly including the NEPAP) would be increased dramatically. Because of the sensitivity of abyssal ecosystems to patterns of POC flux, community structure (including abundance, biomass, species diversity) and ecosystem rates (e.g. organic-carbon mineralization and burial rates) would also be dramatically altered (cf. Smith et al., 1997; Glover et al., 2002). If hypoxic or anoxic conditions were created over large areas of abyssal seafloor, species extinctions would be likely and the nature of sediment geochemical processes (e.g. organic-carbon burial and phosphorous regeneration) would be fundamentally altered (e.g. Canfield, 1994; Van Cappellen & Ingall, 1996). More detailed speculation on the abyssal benthic impacts of ocean fertilization is premature, however, until better resolution of the consequent magnitude and spatial scales of enhanced deep sea POC flux is available.

7.5

Concluding remarks

Because of the low flux of food materials from the euphotic zone, ecosystem structure and function on the NEPAP is largely controlled by annual patterns of POC flux to the seafloor. In particular, the standing crop of NEPAP communities is likely to be very sensitive to long-term changes in export production from the surface ocean. Particles reaching the abyssal seafloor from the euphotic zone appear to be drawn, on an annual basis, from large areas of the surface ocean (in the order of 10 000 km2; Siegel & Deuser, 1997). This fact, plus the low growth and reproduction rates of abyssal benthos (Gage & Tyler, 1991) mean that macro- and megabenthic abundance and biomass integrate surface-ocean production conditions (in particular, export production) over large space and time scales (10 000 km2 and years). Thus, we suggest that the standing stock of NEPAP benthos might be used to monitor large-scale changes in the production regimes of the equatorial Pacific and the oligotrophic central gyre resulting from regional climate cycles, global warming or iron fertilization. In essence, the export-production signal preserved in abyssal macro- and megabenthic biomass may be viewed as having passed through a low-pass filter that smooths short-period (days to months) and small-scale (1–10 km) variability in export production, but responds to broad trends in the flux of labile organic matter to the deep sea. Because macro-benthos, in particular, occur throughout the oxygenated deep sea and are easily sampled using standard quantitative techniques (Gage & Tyler, 1991), macro-benthic standing crop could be readily used to monitor large-scale changes in the ocean’s production cycles as a consequence of global change. Such measurements could augment

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the much more expensive and labor-intensive studies of export production using sediment-trap moorings (e.g. Honjo et al., 1995; Lampitt & Antia, 1997). Because of natural physical stability on the NEPAP, and the small size and fragility of habitat elements and fauna, the NEPAP ecosystem will be very sensitive to even moderate physical disturbance resulting from anthropogenic activities such as manganese-nodule mining. The vastness of the NEPAP habitat may reduce the chances of species extinctions, but commercial-scale mining would almost certainly devastate the biota, and alter geochemical processes, over tens of thousands of square kilometers of seafloor through direct (e.g., nodule removal) and indirect (e.g., sediment redeposition) effects. Recovery of the sediment fauna and geochemical processes following large-scale mining is likely to require decades, while recovery of the nodule biota in mined areas will take millions of years. Substantially more studies of disturbance and recolonization processes, and of biodiversity levels and species ranges, are required before the impacts of nodule mining, and similar large physical disturbances, can be predicted for the NEPAP.

Acknowledgments We would like to thank D.M. Karl and R.R. Bidigare for useful comments on the chapter. This is contribution 6050 from the School of Ocean and Earth Science and Technology, University of Hawaii.

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8

Deep-sea hydrothermal vents and cold seeps Richard J. Léveillé and S. Kim Juniper

8.1 8.1.1

Introduction Deep-sea hydrothermal vents and cold seeps

Deep-sea hydrothermal vents and cold seeps are important components of marine biogeochemical cycles and interactions. Seafloor hydrothermal activity at midocean ridges and in back-arc basins has an important impact on the budgets of heat, water and chemical constituents in the earth-ocean-atmosphere system, for it is primarily in these systems that oceans interact with the lithosphere (Elderfield & Schultz, 1996; Scott, 1997; Herzig & Hannington, 2000). The combination of abundant heat generated by near-seafloor igneous bodies and a nearly unlimited supply of seawater results in widespread alteration of the oceanic crust and formation of abundant metal-rich mineral deposits on and within the seafloor (Scott, 1997). Similarly, cold fluid seeps are sites of heightened interactions between the oceans, sediments and subsurface gases. While seeps do not have the same importance as do vents in terms of metal cycling, they do have a significant role in the carbon cycle of the oceans because of the recycling of organic matter in seep sediments (Herzig & Hannington, 2000). The importance of seafloor discharges to the world’s oceans is made apparent by estimates of fluid flux through vents and seeps. The entire global ocean is thought to be circulated through thermally active seafloor rift zones every 5–11Ma (Herzig & Hannington, 2000), or even less if the ridge flanks are as important as has recently been suggested (see below; Lowell et al., 1995). Estimates of fluid flux at convergent margins suggest that they recycle the volume of water in the oceans every 500 Ma, in a large part via cold seeps (Herzig & Hannington, 2000). 8.1.2

Life at hydrothermal vents and cold seeps

One of the most spectacular features of deep-sea hydrothermal vents and cold seeps is the abundance and novelty of life they support. Vent and seep ecosystems share several fundamental similarities, despite the fact that they are quite different environments that originate from very different geotectonic settings. Hydrothermal vents are characterised by low to high temperature fluids (up to ~400°C), with diffuse or highly focused fluid flow, while seeps are characterised by cold temperature fluids, generally with diffuse, slow fluid flow. Their most

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important common link is the flux of massive amounts of reducing compounds, specifically hydrogen sulphide and methane, from the subsurface. It is these reducing compounds that fuel the abundance of life that flourishes in these environments. Biomass production in hydrothermal vent and cold seep ecosystems is rapid and abundant (Olu et al., 1996; Sarrazin & Juniper, 1999; Tunnicliffe et al., 2003). Life there depends on the chemosynthetic activity of free-living microorganisms (e.g. Karl, 1995) or symbiotic microorganisms living inside animals (e.g. Nelson & Fisher, 1995). At vents, organisms are sustained primarily by mantle-derived volatiles (e.g. H2S, CH4, H2) released by subseafloor inorganic processes, such as volcanic degassing and water-rock interaction. Some of these volatiles may also be produced by subsurface microbial activity (see Section 2.2.2). At seeps, organisms depend largely on methane produced by the degradation of organic matter by thermogenic and biological processes (methanogenesis). In both systems, energy for microbial growth is derived from the oxidation of these dissolved reducing compounds, by exploiting geochemical disequilibria between the fluids and the surrounding seawater or geological materials (Karl, 1995; McCollom & Shock, 1997). The macrofauna, in turn, exploit the abundant microbes either by directly consuming them or by harvesting endo- or ectosymbionts (Nelson & Fisher, 1995; Van Dover, 2000; Tunnicliffe et al., 2003). Although these systems are commonly viewed as being completely independent of photosynthesis, all animals and many microorganisms depend on dissolved oxygen for their metabolism and/or reduced carbon compounds (Karl, 1995; Reysenbach & Shock, 2002; Tunnicliffe et al., 2003). Since molecular oxygen in the oceans is a by-product of photosynthesis in the surface photic zone and on land, there is a critical link between these seafloor systems and the ecosystems in the upper layers (Karl, 1995; Tunnicliffe et al., 2003). At seeps, this link is even stronger as the methane (and H2S) is derived primarily from the degradation of photosynthetically produced organic matter. In both systems, organic compounds from pelagic photosynthesis may also contribute to heterotrophic microbial growth, though the importance of this metabolism is not yet known (Jannasch, 1995; Karl, 1995; Kelley et al., 2002; Tunnicliffe et al., 2003). Microorganisms are the primary agents of energy and chemical transformation in vent and seep ecosystems, and they have been the object of a considerable research effort. Vent microbes have been studied by the isolation and description of new taxa, by physiological characterisation, and more recently by molecular phylogenetic methods. A high microbial diversity at deep-sea hydrothermal vents (e.g. Baross & Deming, 1995; Jeanthon, 2000; Reysenbach & Shock, 2002) is likely the result of the large number of different types of environments and conditions present there (Kelley et al., 2002). Microbial communities at vents are found to exist in all conditions from aerobic to anaerobic, cold to hot (~2–115°C), acidic to slightly basic (Baross & Deming, 1995; Karl, 1995). This high diversity, together with their considerable biomass and potential for

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rapid growth, suggests that they must contribute significantly to global biogeochemical cycling (Reysenbach et al., 2000a). Vent microorganisms are known to influence metal mobility as well as a variety of mineral formation and transformation processes. They are also important in the cycling of C, S and N compounds, and the production of volatiles in the subsurface. However, for many of these organisms we only know their molecular phylogeny based on gene sequencing; relatively few have been isolated and studied in pure culture. Thus, details of the ecology of these microbial systems, the relative importance of autotrophic and heterotrophic metabolism, and the biogeochemical impact of these organisms are poorly known (Karl, 1995; McCollom & Shock, 1997). This is largely due to the inaccessibility of these systems, especially the subsurface portions. The microbiology of deep-sea hydrothermal vents and diversity of microorganisms present there has been reviewed previously (Jannasch & Mottl, 1985; Karl, 1995; Baross & Deming, 1995; Jeanthon, 2000; Reysenbach & Shock, 2002; Kelley et al., 2002) and is beyond the scope of this chapter. Cold seeps also host abundant and diverse microbial communities that have important impacts on geochemical fluxes to the oceans (see review by Aharon, 2000; Orphan et al., 2001). There has been even less isolation and culturing of seep microbes compared to that of hydrothermal vents. Instead, most microbial investigations have used molecular tools to identify those organisms present in seep sediments (e.g. Boetius et al., 2000; Orphan et al., 2001, 2002; Teske et al., 2002). While little is known of their identity, they are clearly involved in methanogenesis, sulphate reduction and oxidation of sulphide and methane, the latter both aerobically and anaerobically (see Section 3.2). Seep microorganisms are also partially responsible for massive carbonate precipitation on the seafloor (see Section 3.3). Although few animal species are shared between vents and seeps, dominant species of both communities (bathymodiolid mussels, vesicomyid clams and vestimentiferan worms) are phylogenetically related and share similar physiology and anatomy (Scott & Fisher, 1995; Peek et al., 1997; Tunnicliffe et al., 2003). Despite these generalities, the ecosystem characters in these habitats are understood only incompletely. For example, little is known of the role of heterotrophy in either of the ecosystems (Tunnicliffe et al., 2003). 8.1.3

Scope of this chapter

Only deep-sea hydrothermal vents and cold-seeps are discussed in this chapter. Shallow vents and seeps are known from a variety of locations, from the littoral zone to several tens of metres (e.g. Holm, 1987; Jensen et al., 1992; Dando et al., 1994a,b, 1995). Shallow vents have many differences from their deeper counterparts. They lack metal-rich and extreme high temperature fluids, as well as large-scale mineral deposits. They also lack typical hydrothermal vent animals. Biomass production in both systems is lower than at deep-sea vents and deep

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cold seeps (by roughly an order of magnitude) and reliance on symbioses between animals and chemolithautotrophic bacteria is minimal, despite the presence of the latter (Dando et al., 1994a, 1995). These shallow systems are also less common, thus likely less important globally. We have included deep-sea cold seeps in this chapter, because of the fundamental similarities with deep-sea hydrothermal vents. In both cases, we have focussed on microbial processes and their impact on geochemical transformations. While interactions between fauna and geochemical processes can be important and diverse in both systems, many of these have been reviewed elsewhere (e.g. Juniper & Sarrazin, 1995; Tunnicliffe et al., 2003). 8.2 8.2.1

Deep-sea hydrothermal vents Distribution and general characteristics

Deep-sea hydrothermal vents are known to occur in all the world’s oceans. Over 200 seafloor hydrothermal sites have been discovered; of these at least 140 show mineralisation, and at least 65 of them are believed to be active (Rona & Scott, 1993; Scott, 1997; see InterRidge internet site [www.intridge. org/ventmap.htm] for exhaustive updated listing). Hydrothermal vents are found at all active mid-ocean ridge and back-arc spreading centres (divergent plate boundaries), as well as at some seamounts. New vents and seafloor hydrothermal systems are continuously being discovered, in part, as a result of improved detection methods and search strategies (Kelley et al., 2002). Much of the 55 000 km-long global ridge system is still poorly explored, though underlying heat sources are not necessarily continuous (Tunnicliffe et al., 2003) and the actual distribution of hydrothermal fields along the global seafloor ridge system is unknown (Lowell et al., 1995). The most spectacular examples of seafloor hydrothermalism are hightemperature black smokers. Beneath these hydrothermal fields, seawater is entrained into the oceanic crust (Fig. 8.1), where it is heated at depths of 2–8 km by the thermal energy of crustal magma chambers or recent intrusive bodies (Alt, 1995; Kelley et al., 2002). With increasing depth, seawater is progressively heated and it reacts with the surrounding rocks, leading to exchange of elements. In the recharge zone, water-rock reactions take place at moderately elevated temperatures (<350°C). These reactions include low-temperature oxidation (e.g. Fe) and alkali fixation, Mg-fixation, anhydrite formation and finally alkali loss from the crust at higher (>150°C) temperatures (Alt, 1995). At greater depths, within the high-temperature reaction zone (~340–465°C), the hydrothermal fluids acquire most of their metals (Alt, 1995; Scott, 1997). In addition, volatiles like H2S, CH4, CO2 and H2 are added to the chemically altered seawater by magmatic degassing, high-temperature leaching of the hot rock and inorganic synthesis at high temperatures (Alt, 1995; Kelley et al.,

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FeO(OH), MnO2, Mn2+, H4SiO4, 3He Bacterial oxidation of CH4, H2S, H2

Metal precipitation of bacterial capsules

PLUME FALLOUT Non-Bouyant Plume

BLACK SMOKER COMPLEX >350oC

PRECIPITATION of SULPHIDES FeS2 ,ZnS, PbS

Hydrothermal Fluid

Metaliferous Sediments

INACTIVE CHIMNEY

115oC THERMOCLINE: ~ Limit of Subsurface Biosphere

LOW TEMPERATURE SEAWATER-BASALT INTERACTION - Microbial Basalt Alteration - Alkali Fixation - Oxidation - Mg-fixation

200oC 400oC R

Decreasing Permeability + Water:Rock

Nontronite Fe-oxyhydroxides

ALTERE

HIGH TEMPERATURE SEAWATER-BASALT INTERACTION - Mg Fixation - Basalt Chloritization - SO42- reduction

WHITE SMOKER >250oC Barite, silica precipitation

MA SS

W AS TI NG

(MICROBIAL ?) SULPHIDE OXIDATION

Low-temperature, off-axis venting (diffuse)

NEUTRALLY BUOYANT PLUME –2o – 40oC

SE AW AT ER EN TR AIN MEN T

Zooplankton

+

-

H , HS , SiO2 , Fe, Mn, Li, Zn, Cu, Pb, Ge, Be, Al, Ba, Rb, CO3

HTRZ

DS

TE WA EA

MAGMATIC DEGASSING

CO2 , H2S, H2 , He, CH4

MAGMA HEAT SOURCE ~1200oC

Fig. 8.1 Diagram illustrating the major geochemical and microbial processes associated with seawater circulation in the crust and hydrothermal activity at mid-ocean ridges. Not all processes or features are found at all hydrothermal vent sites. Diagram not to scale (see text for further details).

2002). This new hydrothermal fluid rises buoyantly to the seafloor and exits the crust at high-temperature, focused vents (i.e. black smokers). Overall, Mg2+ and SO42− are removed from seawater, while H+, Cl−, SiO2, metals and volatiles are removed from the crust and added to the heated seawater. Exiting hydrothermal fluid will have temperatures roughly between 350 and 405°C when no mixing has occurred with cooler seawater (Scott, 1997). Lower temperature vents can result from the dilution of rising hot fluids by entrained crustal seawater prior to exiting at the seafloor. Circulation is ensured at midocean ridges, back-arc basins and seamounts because of the inherent porosity of the lava flows that form the upper crust, as well as by intense faulting and fracturing (Scott, 1997). Since the rate of generation of new oceanic crust is only approximately 5–6 × 1013 kg y−1, the seawater/oceanic crust ratio is very large, which results in extensive seawater–basalt interaction (Scott, 1997; Mottl, 2002). This interaction can also be affected by microbial activity (see Section 2.6). Where undiluted hydrothermal fluids mix with cold seawater at and near the seafloor, minerals precipitate out of solution to form chimney structures and other deposits. These deposits are mineralogically complex, containing sulphides (e.g. pyrite, marcasite, chalcopyrite, sphalerite, wurtzite, galena), sulphates (e.g. anhydrite, barite), silica and oxyhydroxides (e.g. Hannington et al., 1995;

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Scott, 1997). Where these deposits are relatively large, such as at the TAG mound on the Mid-Atlantic Ridge, they resemble ancient volcanic-hosted massive sulphide (VMS) deposits like those presently found in ophiolites (Rona & Scott, 1993; Humphris et al., 1995; Mills, 1995; Goulding et al., 1998; Scott, 1997). Some of the metals at hydrothermal vents are discharged as black smoke along with gases and other dissolved ions, and are expelled into plumes above and away from the vents (see Section 2.4). 8.2.1.1 Geochemical fluxes of gases and elements from hydrothermal vents Hydrothermal circulation is an important mechanism for cooling newly generated oceanic crust; about a third of the earth’s total oceanic heat flux is transferred from the lithosphere to the hydrosphere by hydrothermal flux (e.g. Stein & Stein, 1994; Lowell et al., 1995). Global fluxes of water have been estimated based on estimates of hydrothermal heat flux or the global ocean budgets of Mg and Sr isotopes (German & Angel, 1995; Elderfield & Schultz, 1996; Scott, 1997; Mottl, 2002). Estimates generally range from approximately 1 × 1013 to over 15 × 1013 kg y−1 of water. A large part of the hydrothermal heat flux is believed to occur off-axis, on ridge flanks away from active spreading centres (see Section 2.1.2). In addition, as little as 10% of axial hydrothermal flux may be focused at black smoker-type vents, with the majority being from diffuse on-axis venting (German et al., 1995). Assuming that low-temperature, off-axis heat flow results from the advection of seawater at low to moderate temperatures, an extremely large volume flux (e.g. 2.5 × 1015 kg a−1; Mottl & Wheat, 1994), much greater than that focused through high-temperature vents, would be required to account for the estimated heat flow (German et al., 1995). At such a large volume flux, even small changes in the composition of ridge flank fluids could significantly affect the global hydrothermal flux of chemicals to the oceans (German et al., 1995). Hydrothermal fluids are chemically complex; they contain abundant trace, minor and major elements, as well as abundant volatiles (Elderfield & Schultz, 1996; Scott, 1997; Kelley et al., 2002; Table 8.1). Their composition is controlled primarily by the temperature and composition of the rocks through which they circulate; the composition of the rock can also be influenced by the volume of water that has previously passed through it (Kelley et al., 2002). Compositions can vary markedly from place to place and at a given location over time, though they may be relatively stable over decadal time-scales, at least while the systems are stabilised between volcanic events (German & Angel, 1995; Von Damm, 1995; Herzig & Hannington, 2000). The overall total mass flux from the world’s deep-sea hydrothermal vents is estimated to be on the same order of magnitude as estimates of the flux from the Amazon river, or about 2% of global river discharge (Elderfield & Schultz, 1996; Scott, 1997). Hydrothermal vents can have globally significant mass fluxes that are especially important for their effect on the budgets of elements

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Table 8.1 Comparison of global chemical fluxes from high-temperature (on-axis) hydrothermal vents and rivers

Li K Rb Cs Be Mg Ca Sr Ba Si P B Al Mn Fe Co Cu Zn Ag Pb As Se CO2 CH4 H2 H2 S

Concentration in vent fluids (mol kg−1)

Hydrothermal flux (mol y−1)

River flux (mol y−1)

0.41–1.5 × 10−3 1.35–4.48 × 10−2 9.10–33.0 × 10−6 1.00–2.02 × 10−7 1.00–3.85 × 10−8 – 1.00–8.61 × 10−2 4.80–10.3 × 10−5 8.0–42.6 × 10−6 1.23–2.20 × 10−2 5.0 × 10−7 3.56–5.90 × 10−4 4.0–20 × 10−6 2.50–35.6 × 10−4 7.50–64.7 × 10−4 2.20–22.7 × 10−8 0.97–15.0 × 10−5 4.0–40.0 × 10−5 2.60–3.80 × 10−8 0.90–90.0 × 10−8 0.3–4.52 × 10−7 0.1–7.2 × 10−8 4.2–20 × 10−3 2.5–10 × 10−5 5–100 × 10−5 0.2–1.25 × 10−2

1.52–5.55 × 1010 5.00–16.60 × 1011 3.37–12.2 × 108 3.70–7.47 × 106 3.70–14.2 × 105 −1.60–3.10 × 1012 3.69–31.9 × 1011 1.78–3.81 × 109 2.96–15.8 × 108 4.55–8.14 × 1011 1.85 × 107 1.32–2.18 × 1010 1.48–7.40 × 108 0.930–13.2 × 1010 2.78–23.9 × 1010 8.41–84.0 × 105 3.59–55.5 × 108 1.48–14.8 × 109 0.96–1.41 × 106 3.33–333 × 105 1.11–16.7 × 106 3.70–266 × 104 1.55–7.40 × 1011 0.930–3.70 × 109 1.85–37.0 × 109 0.740–4.63 × 1011

1.4 × 1010 19.0 × 1011 3.7 × 108 4.8 × 106 3.7 × 107 5.3 × 1012 12.0 × 1012 2.2 × 1010 1.0 × 1010 6.4 × 1012 3.3 × 1010 5.4 × 1010 6.0 × 1010 0.49 × 1010 2.3 × 1010 1.1 × 108 5.0 × 109 1.4 × 1010 8.8 × 107 1.5 × 108 7.2 × 108 7.9 × 107 – – – –

Compositional data and river fluxes from Kadko et al. (1994) with modifications based on Edmond et al. (1995) and Scott (1997). Hydrothermal fluxes were calculated assuming a fluid flux of 3.70 × 1013 kg y−1 (Mottl, 2002). Value of magnesium flux from the oceans to the crust is from Mottl (2002).

such as Li, Ca, Mg, Si, Fe and S in the oceans (Table 8.1). They are also important components of the carbon cycle, as they release significant amounts of CO2 and CH4 (and other hydrocarbons) to the oceans. However, accurate estimates of mass flux from hydrothermal vents have been hampered by a lack of knowledge; a key variable being the relative importance of low-temperature venting (Elderfield & Schultz, 1996; Scott, 1997). Also unknown is the global extent and importance of intraplate volcanism and cataclysmic eruptions of hydrothermal fluids that form event plumes (Baker, 1995; Scott, 1997; Kelley et al., 1998). Not all elements dissolved in hydrothermal fluids end up in the ocean. For example, Fe and Mn are generally deposited as sulphides or oxides in chimneys

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and metaliferous sediments near vents, whereas the alkalis are deposited primarily in secondary minerals in altered oceanic crust (Elderfield & Schultz, 1996; Scott, 1997). For most major metals (Cu, Zn, Pb, Co, Cd), the net hydrothermal contribution to seawater composition becomes negligible when compared with the discharge via rivers (German & Angel, 1995). This is due, in part, to the fact that the majority of metals in hydrothermal fluids tend to be removed by precipitation of various sulphide and oxide phases near the seafloor or in hydrothermal plumes (Mottl & McConachy, 1988; Rudnicki & Elderfield, 1993). 8.2.1.2 Off-axis diffuse flow versus axial venting Diffuse, lower-temperature venting from older (several millions or tens of millions of years), tectonised portions of the oceanic crust may be common on the flanks of mid-ocean-ridges, with important implications for heat and chemical fluxes to the oceans (e.g. Stein & Stein, 1994; Wheat & Mottl, 1994; German et al., 1995; Lowell et al., 1995; Mottl et al., 1998; Becker et al., 2000; Wheat & Mottl, 2000). This ridge-flank fluid flow is driven by heat conducted into the crust from the mantle below, as the lithospheric plate cools with age (Mottl & Wheat, 1994). About 70% of global hydrothermal heat flux is believed to occur off-axis in crust older than 1 Ma, despite the lower temperatures of flank venting, because of the much greater surface area (Stein & Stein, 1994; Von Damm, 2001). Circulation through basement rocks on mid-ocean-ridge flanks may be so important that whole-ocean circulation could take place every few hundred thousand years (e.g. Mottl & Wheat, 1994; Elderfield & Schultz, 1996). Because few detailed studies of off-axis venting have been undertaken to date and many have been exclusively near the Juan de Fuca Ridge (JDFR; see references in German et al., 1995; Becker et al., 2000; Wheat & Mottl, 2000), attempts to estimate the overall chemical flux to the ocean through hydrothermal activity and its importance to ocean chemistry, relative to the contribution from rivers, have been hampered by an overall lack of knowledge (Elderfield & Schultz, 1996). Discussion of the relative contributions of off- and on-axis fluxes has centred on comparing the chemical output on the axis to the greater heat (and probably water) flux on flanks (Von Damm, 2001; Mottl, 2002). Recently, Kelley et al. (2001) observed massive carbonate–brucite chimneys associated with low-temperature venting at the Lost City hydrothermal field at 30° N, 15 km away from the main Mid-Atlantic Ridge (MAR) spreading axis on 1.5 Myr old crust. Fluids venting from these structures had temperatures between 40 and 75°C, high pH (9.0–9.8) and compositions indicative of peridotite-dominated (as opposed to basalt-dominated) fluid–rock interaction. Lost City is an example of a previously unknown type of seafloor hydrothermal system that may be much more widespread than the highly localised, magmatically driven hydrothermal vent systems present along mid-ocean-ridge axes; similar systems may be abundant along a significant portion of the global ridge system (Von Damm, 2001; Kelley et al., 2002).

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Subsurface biosphere at mid-ocean ridges

In addition to creating low-temperature seafloor vents, subsurface mixing of oxic and anoxic fluids leads to steep thermal and geochemical gradients (e.g. pH, Eh, metal and gas concentrations) and permits chemical oxidations and aerobic (Karl, 1995; Reysenbach et al., 2000a,b), as well as anaerobic microbial growth below the seafloor at mid-ocean ridges (Table 8.2). The highly fractured and porous subsurface rocks, together with the turbulent mixing of oxidising and reducing fluids, potentially provide abundant microenvironments and interfaces for microbial colonisation and growth. Early reports of microbial activity at 100°C in vent fluids and fluid compositions suggestive of microbial influences, prompted the hypothesis of a subsurface biosphere at deep-sea hydrothermal vents (e.g. Deming & Baross, 1993). Despite the difficulty of direct observation of this potential subsurface biosphere, recent studies have increasingly found barotolerant microbes, thermally stable enzymes and lipids, and culturable hyperthermophilic microbes (e.g. Baross & Deming, 1995; Reysenbach et al., 2000b; Reysenbach & Shock, 2002) in high temperature (e.g. >90°C) vent fluids. In this section, we review the evidence for a subsurface biosphere at mid-ocean ridges and discuss associated biogeochemical processes. 8.2.2.1 Evidence for a subsurface biosphere at deep-sea hydrothermal vents Deming and Baross (1993) proposed the study of accessible black smokers and the ejecta of seafloor diking eruptive events (see below) as indirect windows to the subsurface. Since then, numerous studies have documented a variety of evidence supporting the existence of a subsurface biosphere below mid-ocean ridges, although the quantitative importance of this biosphere remains to be established (Kelley et al., 2002). High concentrations of DNA in vent fluids Higher than background seawater levels of particulate DNA (pDNA) in high-temperature fluids from Endeavour Segment (JDFR), were an early indication of possible subsurface communities (Straube et al., 1990; Deming & Baross, 1993). Deming and Baross (1993) did not find any relationship between total pDNA in vent fluids and percent seawater entrainment, determined by Mg content of vent fluid. Thus, seawater entrainment could not explain the observed concentrations of pDNA in the vent fluids. They did, however, find a positive relationship between pDNA and vent fluid temperature, suggesting that the DNA was produced by thermophilic organisms. “Snowblower Vents”: Post-eruptive venting of sulphur-rich floc Following seafloor volcanic eruptions, abundant matter may be expelled from a type of newly-formed, low-temperature hydrothermal vent (e.g. Haymon et al., 1993;

NH4+, NO2 −

Nitrification

Sulphur reduction – chemoheterotrophy

Sulphur reduction – chemolithotrophic

Organic compounds (e.g. acetate)

Organic compounds H2

CH4 and other C1 compounds*

Methanotrophy/ methylotrophy

Heterotrophy

Fe2+, Mn2+

Metal oxidation

S°, SO42−

S°, S2O32−, SO42−

O2, NO2−

O2

Organic compounds

CO2, organic compounds Organic compounds CO2

CH4, CH3OH, CO, CO2

CO2

CO2

O2, possibly NO3− O2, possibly NO3− O2

H2

CO2

O2, possibly NO3−

HS−, S°, S2O32−, S4O62

Carbon source

Electron acceptor

Electron donor

Hydrogen oxidation

Sulphur oxidation

Pathway

H+ + 4H2 + SO42− → HS− + 4H2O S + H2 → H2S CO2 + SO42− + 6H2 → [CH2O] + S2− + 5H2O CH3COO − + SO42− → 2HCO3− + HS− CH3COOH + SO42 → 2CO2 + S2− + 2H2O 2CH2O + 2H+ + SO42− → H2S + 2CO2 + 2H2O

CH4 + O2 → [CH2O] CH4 + 2H2O → CO2 + 4H2 2CH4 + 2H2O → CH3COOH + 4H2 CH4 + 2O2 → CO2 + 2H2O NH3 + O2 + CO2 → [CH2O] + HNO3

2Fe2+ + 0.5O2 + 2H+ → 2Fe3+ + H2O

S2− + CO2 + O2 + H2O → SO42− + [CH2O] H2S + ½O2 → S° + H2O 2S° + 2H2O + 3O2 → 2SO42− + 4H+ 4H2 + O2 + CO2 → [CH2O] + 3H2O

Reactions

Table 8.2 Known and proposed microbial metabolic pathways at deep-sea hydrothermal vents and cold seeps

Cont

Mesophilic and thermophilic bacteria; hyperthermohilic archaea e.g. Desulfovibrio

Mesophilic and thermophilic bacteria; hyperthermohilic archaea

Mesophilic and thermophilic bacteria

Mesophilic bacteria; Mével et al. (1996)

Mesophilic bacteria detected by activity measurements Mesophilic bacteria; putative Gallionella, Leptothrix Non-thermophilic archaea (Teske et al., 2002)

Mesophilic bacteria (e.g. Beggiatoa)

Examples of organisms and environments

H2

H2 Organic acids

CH4

fatty acids, alcohols, H2

Organic compounds

Methanogenesis

Hydrogen oxidation Iron reduction

Anaerobic methane oxidation (methanotrophy)

Acetogenesis

Fermentation

Organic compounds

CO, fatty acids, alcohols, CO2

CO, fatty acids, alcohols Organic compounds

CO2 Organic acids CH4

CO2, possibly formate, acetate

CO2

NO3− Fe3+ (oxyhydroxides) SO42−

Carbon source

Electron acceptor

CH3COOH → CH4 + CO2

3CH2O + H2O → H3COOH + CO2 + 2H2

CH4 + SO42− → HCO3− + HS− + H2O

4H2 + CO2 → CH4 + 2H2O (CO2 reduction) CH3COOH → CH4 + CO2 (acetate fermentation)

Reactions

Mesophilic and thermophilic bacteria; hyperthermophilic archaea

Identified from molecular analyses Mesophilic bacteria and hyperthermophilic archaea Consortia of methanogenic archaea and SRBs at seeps (e.g. Orphan et al., 2001) and in hydrothermally active sediments (Teske et al., 2002) Limited data from vents

Mesophilic to hyperthermophilic archaea at vents and seeps

Examples of organisms and environments

*C1 compounds are more reduced than CO2 and without C–C bonds. After Karl (1995) and Kelley et al. (2002), with additions.

Electron donor

Pathway

Table 8.2 Continued

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Fig. 8.2 Erupting “snowblower” vent at Axial Volcano (JDFR).

Juniper et al., 1995; Tunnicliffe et al., 1997; Fig. 8.2). Fallouts from these so-called snowblower vents can extend to distances of at least 100 m or more from the vents (Haymon et al., 1993; Delaney et al., 1998). Haymon et al. (1993) first referred to this material as biogenic particulates or biogenic debris, presumably dominated by bacterial mat fragments, from a vent created by, what they interpreted as, hydrothermal flushing of bacterial particles and other biogenic debris from cracks and voids beneath the seafloor at 9° N, East Pacific Rise (EPR). Typically, floc material from these types of vents consists of a dense network of sulphur-rich filaments less than 0.5 μm in diameter (Fig. 8.3) and contains little protein (see Nelson etal., 1989; Wirsen etal., 1993; Juniper etal., 1995). Later work, however, indicated that this material is simply a metabolic end product, consisting primarily of filamentous sulphur, with little biomass produced by sulphur-oxidising bacteria (Taylor & Wirsen, 1997). Wirsen et al. (2002) recently isolated and cultured a coastal marine sulphide-oxidising, chemoautotrophic bacterium that produces hydrophilic filamentous sulphur. They suggested that similar strains might be ecologically important in a number of environments – including hydrothermal vents – contributing significantly to primary production. Despite the relative inorganic composition of this material, its abundance and widespread distribution must be taken as evidence of subsurface microbial production, most likely by sulphide-oxidising bacteria. This material may be produced by in situ populations of bacteria that thrive at depth or by small populations of bacteria that are stimulated by the sudden changes in vent fluid chemistry coupled with minimal disturbance by grazing organisms (Holden et al., 1998). Sustained discharge of these biogenic particles suggests that subsurface microbial production can be continuous and abundant (Haymon et al., 1993; Tunnicliffe et al., 1997).

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Fig. 8.3 Scanning electron microscope photomicrograph of floc material from a “snowblower” – type vent at Axial Volcano, JDFR. The material is dominated by sulphur with little carbon, though it is presumably a microbial product. Scale Bar = 20 μm. Photomicrograph and compositional analysis by C. Levesque.

Culturing of microbial tracers from hydrothermal fluids The presence of anaerobic, thermophilic and hyperthermophilic microorganisms (optimal growth above 45°C and 80°C, respectively) in low-temperature, basalt-hosted diffuse vent fluids expelled after volcanic seafloor eruptions is taken as direct evidence of subsurface origin of the microbes (e.g. Baross & Deming, 1995). If fluids that exit at low temperatures contain high-temperature microbes, those microbes must have been entrained in fluids originating below the surface where temperatures are invariably higher. Thus, isolation and identification of these

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microbes is used as an indicator, or tracer, of a subseafloor ecosystem (Summit & Baross, 2001). Many microbes from deep-sea hydrothermal systems have been isolated directly from venting fluids (see reviews by Baross & Deming, 1995). Some of these isolates can grow at even higher in situ pressures (Reysenbach & Deming, 1991), and some of them possess enzymes that are stable at even higher temperatures (Baross & Deming, 1995; Reysenbach et al., 1998). Pressure stabilization of proteins (e.g. DNA polymerase) has also been cited as a mechanism to extend temperature tolerances of deep-sea thermophiles (Summit et al., 1998). Diverse anaerobic thermophiles and/or hyperthermophiles have been cultured from diffuse low-temperature vents, at a number of locations, at most well-studied ridge systems (e.g., Holden et al., 1998, 2001; Summit & Baross, 1998, 2001; Huber et al., 2002). Huber et al. (2002) suggested that seawaterrelated microorganisms found in vent fluids are not there simply as a result of entrainment of ambient seawater, and that they may in fact be growing and surviving in the subsurface, having adapted to the conditions there. 8.2.2.2 Biogeochemical interactions in subsurface environments McCollom and Shock (1997) proposed that CH4 and H2S are the compounds most likely to be enriched in vent fluids as a result of subsurface microbial processes. Elevated methane concentrations in low and high temperature vent fluids and plumes are often found to be too high to be explained by simple incorporation of mantle-derived methane in end-member fluids (Jannasch & Mottl, 1985; Lilley et al., 1993; Delaney et al., 1998; Kelley et al., 1998; McLaughlinWest et al., 1999). Low values of δ13Cmethane in vent fluids have been found to be within the range of bacterial production by methanogens (Lilley et al., 1993; McLaughlin-West et al., 1999). Most importantly, mesophiles and thermophiles cultured from these sites often include methanogens (Jannasch & Wirsen, 1981; Summit & Baross, 1998; Huber et al., 2002). The isolation of methanogens and their detection by molecular methods point to the importance of these organisms in producing the high concentrations of methane often measured in vent fluids. Lilley et al. (1993) proposed that for the high-temperature fluids, microbial methane may be produced in the lower temperature, downwelling portions of hydrothermal convection cells. Overall, this subsurface production may ultimately enhance the metabolic activities of chemosynthetic organisms at the surface by supplying them with additional H2S and CH4 (McCollom & Shock, 1997; Tunnicliffe et al., 2003). Other types of microbial processes may also be important in the subsurface (Table 8.2). Microorganisms, such as sulphate-reducers have also been cultured from low-temperature fluids, and cell morphologies indicative of sulphide-, ironand methane-oxidizing microorganisms (and protozoa) have been observed in vent fluids (e.g. Holden et al., 1998). Heterotrophs are also present at hydrothermal vents, though little is known about their abundance and distribution, or

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of their role in carbon cycling at vents. One source of organic matter for heterotrophs in the subsurface may be sediments buried at an earlier stage of ridge evolution or accumulation within the crust of organic matter from previous microbial oxidative metabolism (Delaney et al., 1998). Depending on the rate of supply of hydrothermal nutrients to the extrusive layer between eruptive diking events, slowly growing or dormant oxidative mesophiles could survive within the upper crust, providing both the source of organic carbon for heterotrophs and the seed organisms for the massive microbial blooms that follow seafloor eruptions. 8.2.3

Seafloor microbe-mineral interactions at hydrothermal vents

The interaction of hot, metal-rich vent fluids and cold seawater creates a variety of conditions that can potentially support diverse microbial life (Table 8.2). The oxidation of dissolved hydrogen sulphide coupled with CO2 fixation is likely the most important component of carbon and energy flow at deep-sea vents. H2S is usually the most abundant chemosynthetic energy source in hydrothermal fluids, and its oxidation provides the most energy of the metabolisms used at vents (Karl, 1995; Baross & Deming, 1995; McCollom & Shock, 1997; Fig. 8.4). Other chemoautotrophic reactions (oxidation of methane, Fe2+, Mn2+; reduction of sulphate/sulphur) provide less energy (Fig. 8.4), but may still be important in hydrothermal systems, particularly in terms of their effect on geochemical fluxes (McCollom & Shock, 1997). 8.2.3.1 Microbial distribution and activity at vents Sulphide-oxidising bacteria abound at vents, they are usually mesophilic, and occur as free-living (see review by Juniper & Tebo, 1995) and symbiotic forms (Nelson & Fisher, 1995). Mesophilic sulphur and iron oxidisers are likely favoured by the presence of reduced sulphur and iron at vents, and often produce conspicuous filamentous mats (Emerson & Moyer, 1997; Reysenbach & Cady, 2001; Emerson & Moyer, 2002). However, many presumed occurrences of metaloxidising bacteria (iron, manganese) are only based on microscopic studies (see Juniper & Tebo, 1995). Emerson and Moyer (2002) recently isolated an obligate, microaerophilic Fe oxidiser from mats at Loihi Seamount, where fluids are rich in Fe, though poor in H2S. Methane- and hydrogen-oxidising bacteria have also been found in hydrothermal plumes (see below) and detected by gene sequencing of vent samples, although they have rarely been studied in isolated culture (Kelley et al., 2002). Progress on the isolation, identification and laboratory-based physiological and taxonomic investigations of selected vent microbes has been remarkable, but substantive ecological investigations have lagged behind (Karl, 1995; Reysenbach & Shock, 2002). The quantitative role of chemolithautotrophy as a primary source of carbon and energy at vents remains unknown and the role of

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1000

Sulphide/sulphur oxidation

cal/kg vent fluid

Fe2+ oxidation to hematite 100 Methanotrophy S reduction

Fe2+ oxidation to Fe3+ Mn2+ oxidation 10

Sulphate reduction Methanogenesis

25

50

75

100

125

Temperature °C Fig. 8.4 Theoretical estimates of energy available from hydrothermal vent fluids for various chemosynthetic reactions used by vent microbes. Oxidative reactions are shown as dashed lines and reductive reactions as solid lines. Iron reduction provides too little energy to appear on plot. H2 oxidation was not calculated. From McCollom and Shock (1997).

heterotrophy is unquantified (Karl, 1995; Tunnicliffe et al., 2003), and there are few detailed studies of chemosynthetic microbial activity at vents (e.g. Wirsen et al., 1993; Eberhard et al., 1995). McCollom and Shock (1997) calculated that microbial chemoautotrophic productivity at hydrothermal vents may be in the order of 1013 g biomass y−1, which is equal to approximately 0.02% of the global photosynthetic primary productivity in the world’s oceans; this chemosynthetic productivity at vents is likely significant on a local scale (McCollom & Shock, 1997). The aerobic and anaerobic heterotrophic microbes present at vents often have complex nutrient requirements (e.g. organic acids, carbohydrates, elemental sulphur); some are believed to be mixotrophic, using both autotrophic and heterotrophic metabolisms (Karl, 1995; Kelley et al., 2002). Organic matter in sulphide mineral chimneys may come from other dead and/or pyrolysed microbes and invertebrates (e.g. worm tubes and muccus) during growth of chimneys. At most hydrothermal sites, thermal environments at moderate temperatures (50–90°C) are

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dominated by Eubacteria, while high temperature environments (>90°C) are dominated by Archaea (Reysenbach & Shock, 2002). While much microbiological attention has focused on thermophiles, lower temperature microorganisms (some belonging to nutritional groups other than S-oxidizers) may, in fact, be the most abundant in vent environments, though very little is known of their taxonomic or ecological significance (e.g. Baross & Deming, 1995). Microbes in hydrothermal sulphide chimneys In sulphide chimney structures and underlying sulphide mounds, steep environmental gradients of temperature, pH, redox potential and chemistry can be formed in equilibration between vent water and seawater, and these provide diverse microhabitats and inorganic energy sources for microbial communities (McCollom & Shock, 1997; Delaney et al., 2001; Takai et al., 2001; Fig. 8.5). The exterior surfaces of high-temperature mounds and chimneys are frequently encrusted with an abundance of microbes, often filamentous and metal-depositing bacteria (e.g. Baross & Deming, 1985; Juniper & Fouquet, 1988), as well as thermophiles (Baross & Deming, 1995). Typically, there is a decrease in biomass (based on direct counts or lipid analysis) from the outer surface (~107–1010 cell g−1 dry weight) to a few centimetres towards inner regions (~105–106 cells g−1 dry weight) to below detection (i.e. ~103–104 cells g−1) for the inner most active portions (Chevaldonné & Godfroy, 1997; Harmsen et al., 1997; Delaney et al., 2001; Takai et al., 2001). Lipid OXIC

ANOXIC

SEAWATER 2°C, pH~8

Organics, SO42–, O2

Fe-oxides

MESOPHILIC S-Fe-Mn Oxidizers Methanotrophs Heterotrophs

H2S, CH4, H2, Fe2+, Mn2+,

THERMOPHILIC H2 oxidizers Methanogens Heterotrophs S-Reducers

? Microbial Abundance

100° C

200° C 300° C

VENT FLUID 350° C, pH ~3-5

Fig. 8.5 Cross-sectional diagram of a black-smoker chimney illustrating hypothetical microbial habitats. These habitats may shift and change as fluid flow varies or stops. Modified after McCollom and Shock (1997).

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analysis and molecular phylogeny suggest that there is a prevalence of Archaea in the inner parts of the chimneys, while both Bacteria and Eukaryotes are more common on (or near) the surfaces (Hedrick et al., 1992; Chevaldonné, 1996; Delaney et al., 2001). This suggests that these groups of microbes have adapted to different niches, likely, as a result of their different tolerances to environmental factors. Microbial diversity in chimneys may be high, as suggested by novel cell morphologies, presence of mixed communities and development of biofilms, as viewed by scanning electron microscopy (SEM; Chevaldoné, 1996; Harmsen et al., 1997; Delaney et al., 2001), and more recently by phylogenetic studies (Harmsen et al., 1997; Delaney et al., 2001; Takai et al., 2001). Community structure (diversity and density) in chimneys can vary on a scale of a few centimetres and likely reflects the formation of environmental gradients of temperature, pH, redox potential and chemistry (Takai et al., 2001). Wirsen et al. (1993) measured high chemosynthetic biomass production (14CO2 fixation and RuBisCo activity) on the surfaces of sulphide deposits at TAG and Snake Pit (MAR) under mesophilic conditions (20°C), but weak inside the deposits under thermophilic conditions (80°C), decreasing to below detection within 1.5 cm. Muyzer et al. (1995) observed the sulphate reducer Desulfovibrio amongst other dominant phylotypes (including sulphur-oxidizers) from Snake Pit/MAR. They suggested that sulphate-reducing bacteria may serve as a source of sulphide and complement sulphur- and sulphide-oxidising bacteria, and are thus very important ecologically in the hydrothermal vent ecosystem. A closely associated consortium of chemosynthetic sulphate-reducers and sulphide-oxidisers along geochemical gradients has been proposed to exist in the walls of hydrothermal vent chimneys (McCollom & Shock, 1997). Despite these studies, the overall microbial ecology of sulphide deposits is not well known. Optimal growth may occur at oxic/anoxic interfaces, as H2S and Fe2+ are rapidly oxidized in seawater (Karl, 1995). While much effort has concentrated on isolating, culturing and identifying thermophiles and hyperthermophiles, especially using molecular methods, the physiology and biochemistry of these organisms are poorly understood. Furthermore, most work has focused on actively venting sites, with little work on extinct or inactive chimneys and seafloor deposits. The occurrence, abundance and distribution of the communities associated with the formation of diverse geochemical and physical gradients remains to be elucidated completely (Takai et al., 2001; Kelley et al., 2002). Inactive chimneys, hydrothermal and sulphide sediments and mounds The interior of sulphide mounds has been likened to an exposed mesoscale view of the subsurface environment, and hence is of great interest to microbiologists (Deming & Baross, 1993). This environment may also provide a continuum of bacterial populations between high and low temperature hydrothermal

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systems, thus representing an important transitional stage in the hydrothermal ecosystem (Severmann et al., 2000). Although sampling such environments by drilling is difficult and aseptic methods are never ideal, reasonable precautions have been shown to be sufficient to avoid substantial contamination (Reysenbach et al., l998). Reysenbach et al. (1998) found microbes in the first metre of drill core from TAG (ODP Leg 158) and were able to produce enrichment cultures of hyperthermophilic, S-reducing organisms, but they could not rule out that these organisms were entrained by fluids. Samples of deeper (i.e. 1–52 m), hotter parts of drill core produced no microbial growth and cell counts were too low to be reliable. Microbial communities were also found to exist in sulphide sediments from inactive parts of TAG, and they are presumed to be involved in cycling of S, Fe and Mn there (Severmann et al., 2000). Cragg et al. (2000) detected significant cell numbers in hydrothermal sediments from Middle Valley (JDFR) to depths of at least 30 mbsf. They found that bacterial populations decreased with depth and became undetectable at 110°C, though intact cells were observed at depths corresponding to ~155–185°C. Their distribution thus seems to be influenced primarily by temperature, though little is known of their ecology and physiology. Sulphate reduction is common in hydrothermally active sediments, such as in Guaymas Basin, even at temperatures up to 110°C (Jørgensen et al., 1992). A phospholipid fatty acid study of hydrothermally influenced sediments from Middle Valley (JDFR), found eubacterial membrane lipids, but no Archaeal lipids at temperatures of 60–130°C (Summit et al., 2000). 8.2.3.2 Biomineralisation at vents The high flux of dissolved elements and the prolific mineralisation that takes place around vents, combined with the abundance and diversity of microorganisms, creates a very high potential for biomineralisation (Juniper & Tebo, 1995). Numerous studies have implicated vent microorganisms in the formation of various minerals, though few have provided convincing biological and geochemical evidence. Jannasch and Wirsen (1981) cautioned against indiscriminately attributing metal accumulation on vent bacteria to direct or indirect effects of the organisms themselves, because non-specific metal deposition is common around vents, especially associated with organic-rich coatings that may contain microorganisms (Léveillé & Juniper, 2002). Fe-Mn oxides/hydroxides Filamentous, putative microbial textures, often with morphologies resembling iron-oxidising bacteria, are extremely common in Fe-Mn-Si oxyhydroxides from a variety of mid-ocean ridges (e.g. Juniper & Fouquet, 1988; Bogdanov et al., 1997; Fortin et al., 1998; Boyd & Scott, 1999; Halbach et al., 2001; Kennedy et al., in press; Léveillé & Juniper, 2002), seamounts (Alt, 1988; Bogdanov et al., 1997; Boyd & Scott, 1999; Boyd &

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Scott, 2001) and shallow hydrothermal vents (e.g. Holm, 1987). Similar filamentous textures have also been found in seafloor stockwork, as well as in ancient hydrothermal Fe-rich cherts and marine VMS deposits of various geological ages (Juniper & Tebo, 1995; Konhauser, 2000), suggesting that bacteria have had an important role in Fe-biomineralisation at hydrothermal vents through time (Konhauser, 2000; Humphris et al., 2002). Poorly crystalline Fe-hydroxides are frequently cemented by later-stage silica, thus ensuring good preservation of the deposits and their textures (Juniper & Fouquet, 1988; Juniper & Tebo, 1995). Microbial enhancement of Fe-oxide precipitation on seamounts in the eastern Pacific (Alt, 1988) and diagenetic alteration of secondary mineralisation of sulphide debris (Goulding et al., 1998) have also been inferred to contribute to formation of ochre and Fe-oxide muds. In many cases, the deposits consist almost entirely of long, delicate filaments, remarkably similar in morphology to the genera of iron oxidising bacteria (e.g. Gallionella, Leptothrix, Leptospirillum, Metallogenium; Alt, 1988; Boyd & Scott, 2001; Kennedy et al., in press). At Loihi Seamount, CO2- and Fe-rich, sulphide-depleted fluids diffuse up through extensive microbial mats (stained orange with iron oxide coatings), resulting in the precipitation of Fe-oxide on bacterial filaments (Karl et al., 1988). Although no H2S was detected in Loihi vent fluids, deposits of elemental sulphur suggest that oxidation of vent water H2S may be occurring. Emerson and Moyer (2002) reported the isolation of an obligate, microaerophilic iron oxidiser from Loihi mats that were encrusted in Fe-oxides. They suggest that Fe-oxidising bacteria play a major role in Fe oxidation at Loihi and many hydrothermal vents around the world. True autotrophic iron bacteria use the energy from Fe2+ oxidation to Fe3+ in order to fix CO2, a metabolism which can likely lead to Fe-oxide precipitation (Emerson, 2000; Kennedy et al., in press). Non-Fe-oxidising bacteria can precipitate Fe indirectly by locally altering pH/Eh and/or by adsorbing Fe3+ onto their cell surfaces. Bacterial mats and/or bacterial strands are also preferred sites for primary silica precipitation because of their high surface area and a number of potential precipitation sites (Konhauser, 2000; Humphris et al., 2002). Extracellular polymers can also non-specifically adsorb positively charged Fe-hydroxides (Humphris et al., 2002). Baross and Deming (1985) provided evidence that at an active black smoker from the East Pacific Rise, iron was deposited in significantly greater amounts on mineral surfaces colonised by microorganisms (e.g. microbial mats) than on uncolonised surfaces. Thick mats may also promote conductive cooling of silica-rich upwelling fluids, leading to precipitation of amorphous silica on and within bacterial strands. Manganese, because of its relative stability in seawater (Ehrlich, 1990), is a potentially important energy source associated with hydrothermal vents. The high flux of reduced Mn2+ to neutral to slightly alkaline oxygenated waters and the relatively low amount of dissolved organic carbon, make hydrothermal

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vents excellent systems to find Mn autotrophic or mixotrophic organisms (Juniper & Tebo, 1995). A variety of putative Mn-oxidising bacteria have been observed by electron microscopy in samples collected from hydrothermal vent environments and plumes (see references in Juniper & Tebo, 1995; Section 2.5). A number of Mn-oxidising bacteria have also been isolated from hydrothermal vents (Ehrlich, 1990; Juniper & Tebo, 1995). However, because these microbes are associated with low temperature diffuse venting, they have been poorly studied. A combination of Mn-encrusted bacteria-like particles and greatly accelerated precipitation rates of Mn-oxides on experimental substrates at Cross Seamount were taken as strongly suggestive of biological mediation of Mn-oxide precipitation (Bertram & Cowen, 2000). Some vent microbial mat bacteria exhibit Mn2+ oxidising activity coupled with ATP synthesis; in some cases the Mn2+ oxidising activity is associated with cell envelopes (Ehrlich, 1990). Layer-silicates Recent studies have also demonstrated the potential microbial influence on clay mineral (layer silicates) formation at hydrothermal vents. Bacterial cells covered (or completely replaced) with a Fe-rich silicate mineral (putative nontronite), in some cases oriented in extracellular polymers (as revealed by TEM analysis), were found in deep-sea sediments of Iheya Basin, Okinawa Trough (Ueshima & Tazaki, 2001), and in soft sediments, and on mineral surfaces in low-temperature (2–50°C) waters near vents at Southern Explorer Ridge in the northeast Pacific (Fortin et al., 1998; Fig. 8.6). The Fe-silicate is believed to form as a result of the binding and concentration of soluble Si and Fe species to reactive sites (e.g. carboxyl, phosphoryl) on EPS (Ueshima & Tazaki, 2001). Formation of Fe-silicate may also involve complex binding mechanisms, whereas metal ions such as Fe possibly bridge reactive sites within cell walls to silicate anions to initiate silicate nucleation (Fortin et al., 1998). Alt (1988) also reported the presence of nontronite associated with Mn- and Fe-oxide-rich deposits from seamounts on the EPR. The presence of bacteria-like filaments within one nontronite sample was taken to indicate that bacterial activity may have been associated with nontronite formation. Although the formation of clay minerals at deep-sea hydrothermal vents has not received much attention, it seems probable that based on these studies, biomineralisation of clay minerals is ubiquitous in these environments. Sulphur mineralisation Sulphur is ubiquitous in minerals at hydrothermal vents (e.g. native sulphur, metal sulphides, sulphates), and the formation of many of these minerals is potentially influenced by the associated microbial activity at vents. For example, microbial filamentous sulphur formation has recently been proposed as a widespread activity in mixed H2S-O2 environments and may be a quantitatively important microbial process in hydrothermal vent

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1 μm

Fig. 8.6 Transmission electron microscope photomicrograph of Fe-rich layer silicate surrounding microbial remains, which appear as “holes” in this unstained section. Sample from Southern Explorer Ridge. The mineral has a composition similar to nontronite. From D. Fortin et al., Amer. Mineral., 83, 1399–1408. Reproduced with permission. Copyright Mineralogical Society of America (1998).

systems (Taylor & Wirsen, 1997; Taylor et al., 1999). Microbial mats may also provide surfaces for precipitation of sulphide minerals (Juniper & Fouquet, 1988; Zierenberg & Schiffman, 1990). Surfaces of hydrothermal crusts of highly altered basalt from the Sea Cliff hydrothermal field, northern Gorda Ridge, were previously colonised by bacterial mats which were locally preserved by replacement and overgrowth of bacterial filaments by metal sulphide minerals and amorphous silica (Zierenberg & Schiffman, 1990). The bacterial filaments within the mats were selectively replaced by chalcopyrite (CuFeS2), Ag-As-rich sulphide (prousite, pearceite) and rarely galena (PbS). Zierenberg and Schiffman (1990) suggested that bacterially mediated processes selectively precipitated silver, arsenic and copper, and that biological processes may contribute to precious-metal enrichment in some seafloor hydrothermal base-metal sulphide deposits. Cook and Stakes (1995) described abundant fossilised tube structures (up to ~1 cm wide) from large black smoker massive Fe-Zn-sulphide edifices sampled by a short (~30 cm) drill core. Concentric layering within the tube walls was proposed to be initiated by the presence of worms and possibly

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controlled by microbes before inorganic reactions dominated the final stages of development. They suggested that bacteria may be responsible for the transition from sulphate to sulphide assemblages by providing a nucleation site for sulphide accumulation. These studies clearly demonstrate that the abundant sulphideoxidising and sulphate-reducing bacteria at deep-sea hydrothermal vents may have an important influence on the mineralisation of sulphur-bearing minerals in these environments. 8.2.3.3 Fossilisation of microbes at vents The propensity of bacterial surfaces to serve as nucleation sites for mineral precipitation often leads to the preservation or fossilisation of bacterial remains within mineral deposits. In fact, the study of biomineralisation processes at hydrothermal vents often begins with the observation of microbe-like fossils in mineral samples. Fossilised hydrothermal vent microbes have been identified in deep sea VMS deposits as old as 3.235 Ga at Sulfur Springs, Australia (Rasmussen, 2000). The combination of rapid mineral precipitation and extensive microbial growth associated with vents leads to a high potential for preservation of microbial remains (Juniper & Fouquet, 1988; Juniper et al., 1995). Precipitation or binding of iron to cell surfaces may contribute to their preservation prior to silicification (Ferris et al., 1988; Reysenbach & Cady, 2001). Filamentous microfossils in silica-rich and Fe-oxide deposits are abundant in modern and ancient hydrothermal deposits (e.g. Juniper & Fouquet, 1988; Little & Thorseth, 2002). Al-Hanbali et al. (2001) provided microscopic and chemical evidence that cell-like structures found in hydrothermally precipitated, Fe-rich cherts (silica) at TAG are recently silicified microorganisms. The criteria they used for assessing biogenicity included the morphological similarity of the structures to living bacteria (size, form and arrangement as viewed by SEM) and chemical signatures that resemble biological composition (elevated C and N concentrations and N/C ratios consistent with those of marine bacteria). In comparison, Rasmussen (2000) reported pyritic putative microfossil filaments encased in hydrothermal Archaean cherts. Biogenicity was inferred from their sinuous morphology, length-wise uniformity and intertwined habit, which are comparable to those of Archean and younger microfossils (e.g. Juniper & Fouquet, 1988). 8.2.3.4 Bacterial weathering of sulphides Although less energetically efficient than the oxidation of hydrogen sulphide (Fig. 8.4), the oxidation of metal sulphides could potentially support chemosynthesis at seafloor massive sulphide deposits long after hydrothermal activity had ceased, even in well-buffered seawater (Eberhard et al., 1995; Juniper & Tebo, 1995). Newly formed sulphide deposits are rapidly subjected to oxidation upon contact with ambient seawater, and some show micro-scale weathering features (e.g. etch pits on mineral surfaces; Verati et al., 1999),

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large-scale orange-brownish coatings (especially on inactive spires/chimneys) consisting mainly of secondary Fe-oxyhydroxide minerals (see Juniper & Tebo, 1995; Goulding et al., 1998), and evidence of talus and mass wasting (e.g. Scott, 1997). Despite these observations, the mechanisms involved in the weathering of seafloor sulphides (either geochemical or microbial) remain poorly understood (Juniper & Tebo, 1995). In their study of bacterial mats on polymetal sulphide deposits at the TAG and Snake Pit sites on the MAR, Wirsen et al. (1993) and Eberhard et al. (1995) isolated obligate chemoautotrophic, non-acidophilic, S-oxidising bacteria that could oxidise sterilised sulphides at near-neutral pH. Eberhard et al. (1995) also assayed S-oxidation of sulphides by aerobic, mesophilic S-oxidising bacteria via in situ, shipboard and laboratory experiments. In all cases, substantial fixation of 14CO2 occurred at near-neutral pH by natural populations of bacteria covering sulphide minerals. Avery et al. (2000) counted a higher than normal density of microbes on vent sulphide minerals with associated pits and irregular grain boundaries on the sulphide minerals. It has been often postulated that non-acidophilic, Fe-oxidizing bacteria (e.g. Gallionella, Leptothrix) are active at vents, deriving their metabolic energy from reduced ferrous iron (Wirsen et al., 1993; Juniper & Tebo, 1995; Verati et al., 1999). However, these studies have relied mainly on microscopic studies of morphology alone. Obligate, microaerophilic Fe-oxidizers have been enriched and isolated from low temperature vent fluids at Loihi Seamount, where microbial mats are typically encrusted with Fe-oxides (Emerson & Moyer, 2002). Chen and Chen (2001) identified minor amounts of nitrate-dependent ferrous iron oxidizers by 16S rRNA oligonucleotide probes in an extremely anoxygenic environment in the Okinawa Trough, off the east coast of Taiwan. These studies demonstrate that bacteria are likely exploiting the vast energy available in hydrothermal vent sulphides, and these deposits could potentially serve as long-term energy sources for the chemolithoautotrophic production of microbial biomass on the seafloor. However, the quantitative significance of this energy source and the environmental importance of this form of metabolism remains to be determined. 8.2.4

Biogeochemical interactions in hydrothermal plumes

8.2.4.1 General features of hydrothermal plumes Hydrothermal plumes are the water column manifestation of mid-ocean ridge hydrothermal activity (e.g. Baker et al., 1995). They form above areas of venting as a result of the buoyancy of the heated fluids; as the fluids rise, they entrain ambient seawater, thus continuously increasing the plume volume until neutral buoyancy is achieved and the plume disperses laterally (Baker et al., 1995). These warm, buoyant plumes rise up to several hundred metres above the seafloor and extend for several kilometres away from vents [e.g. see reviews by Baker

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et al. (1995); Lupton (1995) and Baker (1998)]. In addition to their high heat content, plumes contain significant concentrations of volatiles (e.g. H2; CO2; CH4; Kelley et al., 1998; McLaughlin-West et al., 1999), metalliferous (e.g. Fe, Mn) and non-metalliferous elements (e.g. sulphur; Feely et al., 1999; Khripounoff et al., 2001), and organic matter (e.g. Baker & Massoth, 1988). Chronic plumes are the overall result of hydrothermal discharge (diffuse or focused), and they overlie all active hydrothermal areas on ridge crests worldwide (e.g. Baker, 1995). In contrast, event plumes form episodically as the result of massive volcanic eruptions or dyking eruptive events (Baker, 1995). The latter are relatively uncommon, though they may be equivalent to many months of chronic plume formation (Baker, 1995). In either case, plumes turbulently mix with ambient seawater and, when they achieve neutral buoyancy, they represent an approximately 104–105 dilution of vent fluids with seawater (Baker et al., 1995). 8.2.4.2 Microbial ecology of hydrothermal plumes Massive bacterial output from fissures and vents created during diking eruptive events is believed to characterise the early stages of volcanic events and formation of hydrothermal plumes (Karl, 1995). Bacterial (and viral) biomass is elevated in buoyant and non-buoyant plumes, not only because of the entrainment of microbes derived from vent environments, but also of enhanced microbial activity and growth within the plume (Haymon et al., 1993; Lilley et al., 1995; Winn et al., 1995; Holden et al., 1998; Juniper et al., 1998; Summit & Baross, 1998; Cowen et al., 1999). These studies have shown that typical abundances of bacteria in plumes can be 2–3 times as high as background seawater values and that these abundances often correlate with temperature. The abundant volatiles and metals released into plumes are likely to be used by a variety of microorganisms as energy sources (Delaney et al., 1998; McCollom, 2000). Because event plumes represent the sudden injection of exploitable reducing chemical substrates, as well as inhibitory constituents, they are likely to induce successional changes in the microbial community structure and activity within plume waters over time (Cowen et al., 1998). For example, in studies following the 1998 Axial Volcano eruption, abundant putative bacterial sulphur filaments were observed in August 1998 (Feely et al., 1999), though they were not initially found in plumes in February 1998 (Cowen et al., 1999). The origin of microbes present in hydrothermal plumes is believed to be either from a subsurface biosphere maintained by the rapid mixing processes that take place during plume formation or from entrained seawater microbes (Winn et al., 1995; Cowen et al., 1998, 1999). In the latter case, most of the organisms would not be able to grow, but a small segment of the population may thrive and multiply as plumes move away from the ridge crest. Similar forms to plume bacteria are common on rock surfaces near vents and on sulphide structures (Jannasch &

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Wirsen, 1981; Juniper & Tebo, 1995). To date, however, little phylogenetic or culture work has been done specifically on plume microbes. 8.2.4.3 Microbial productivity and organic carbon in plumes The amount of particulate organic carbon (POC) in hydrothermal plumes may be large relative to that in abyssal waters, and hydrothermal plumes may thus have an important influence on the organic C cycle in the deep sea, especially locally (i.e. scale of kilometres or tens of kilometres; Winn et al., 1995). Organic carbon in plumes is derived from a variety of sources, including subsurface microbial and thermo-chemical processes, export from abundant biological communities associated with active vent fields, background deepwater compounds entrained during plume formation, descending particles ultimately derived from surface water photosynthesis, and in situ microbial production (Cowen et al., 2001 and references therein). Additional mid-depth inputs are supplied by presumed chemolithotrophic-based carbon production and from mid-water zooplankton (Cowen et al., 2001 and references therein). In situ microbial production is potentially the most important source of organic C to buoyant and especially neutrally buoyant plumes (Lilley et al., 1995; Winn et al., 1995; Cowen et al., 2001). However, plume productivity is poorly constrained; only limited simultaneous measurements of substrate oxidation rates and CO2 incorporation rates have been made in hydrothermal plumes for key substrates. For example, rates of dark CO2 assimilation and methane oxidation in plumes from the Manus and Lau basins were found to be relatively high (1200–2500 ng C−1 day−1 and 1300 ng C−1 day−1, respectively), but dropped considerably with increasing distance from the vents (Lein et al., 1997). Although up to 60% of the global mid-ocean ridge system may be overlain by hydrothermal plumes, the global hydrothermal production of organic carbon is probably less than about 1% of the total oceanic photosynthetic production (Winn et al., 1995). However, most hydrothermally driven productivity occurs likely within a ~10 km wide swath over the ridge system. Along this productivity corridor, organic carbon produced from hydrothermal energy sources could be ~5 times greater than the downward flux of photosynthetically derived carbon at plume depths (Cowen & German, 2002). Particulate organic carbon could also be a significant source of energy for heterotrophic microbial growth in plumes (Winn et al., 1995). However, vertical particle flux data are scarce and often variable: ~2–4mg C m−2 d−1 for descending particles at Endeavour, where a clear hydrothermal signal (Fe and Mn) is observed (Roth & Dymond, 1989; Cowen et al., 2001; Bertram et al., in press), ~10–25 mg m−2 d−1 at the Rainbow, MAR site where a smaller but distinct hydrothermal flux is present, consistent with settling of far-field hydrothermal Fe-oxyhydroxides. Thus, the mid-ocean ridge corridor could be a 55 000 km long source of labile organic carbon that is independent of seasonal upper ocean production

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cycles, and that which could support deep-water zooplankton, both vent and non-vent varieties (Karl, 1995; Cowen & German, 2002). Bio-acoustic surveys and zooplankton net-tows have documented anomalously high concentrations of zooplankton (10–100 times above the background) in the vicinity (especially within a 100 m layer above the plumes) of hydrothermal vents (Burd & Thomson, 1994; Cowen et al., 2001; Wakeham et al., 2001), and the zooplankton density tends to decrease with the distance from active venting as a result of nutrient enrichment around the hydrothermal area and the abundance of free living bacteria (Khripounoff et al., 2001). First order calculations of organic carbon input (5–22mg C m−2d−1) required to sustain the observed epi-plume zooplankton anomalies at Endeavour Segment are comparable both to measured POC flux to pi-plume depths (2–5 mg C m−2 d−1) and to estimates of the total potential in situ organic carbon production (2–9 mg C m−2 d−1) from microbial oxidation of H2, CH4 and NH4+ in the hydrothermal plume (Cowen et al., 2001). 8.2.4.4 Biogeochemical interactions in plumes The massive output of microbial biomass that characterises the early stages of event plumes potentially affects, and is affected by, hydrothermal constituents like H2S, Mn2+, Fe2+, H2, CH4. These reducing substances represent important energy sources as in other hydrothermal environments. Elevated methane and hydrogen concentrations in plumes (10–1000 times background seawater levels) are often noted following seafloor eruptions (e.g. de Angelis etal., 1993; Kelley et al., 1998; McLaughlin-West et al., 1999). Stable carbon isotopic compositions of methane indicate numerous sources in mid-ocean ridge systems, including thermogenic decomposition of organic matter and microbial methanogenesis (likely in the subsurface), though the latter cannot be easily shown unequivocally (Lilley et al., 1993, 1995; McLaughlin-West et al., 1999). Hydrogen can also be produced by a number of redox reactions within hydrothermal systems, including the reaction of ferrous minerals and water (McLaughlin-West et al., 1999). In some cases, hydrogen concentrations decrease rapidly over subsequent weeks and months, and they do not correlate well with other plume signals (i.e. CH4, 3He, temperature) suggesting that H2 concentrations decrease by a mechanism other than dilution, likely microbial oxidation (McLaughlin-West et al., 1999). In contrast, methane and Mn2+ are somewhat longer-lived in plumes, though microbial activity related to the oxidation of both has also been demonstrated (Lilley et al., 1995 and references therein). Measurement of microbial oxidation rates of methane in plumes at Endeavour gave average turn-over times of about 2 weeks or longer (de Angelis et al., 1993), which is generally longer than turn-over times for hydrogen (McLaughlin-West et al., 1999). However, microbial oxidation rates of methane following the 1998 Axial Volcano eruption were found to be negligible (McLaughlin-West et al., 1999).

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Mn oxidation and scavenging in plumes appear to be dominated by microbial activity (see review in Winn et al., 1995). Plume bacteria are often characterised by capsules of abundant extracellular polymer matrices that are believed to scavenge Mn from solution (e.g. Cowen et al., 1986, 1999). Radiotracer experiments (54Mn uptake) and elevated microbial biomass at plume depths suggest that microbial metabolic activity enhances Mn scavenging (Cowen et al., 1986). However, it is not clear if this activity includes active Mn2+ oxidation or is simply due to non-enzymatic processes. In comparison, Fe oxidation and deposition appear to be much less common in plumes. About half the Fe in the hydrothermal fluids combines with H2S and is rapidly transformed into Fe sulphides within a few seconds of release (e.g. Rudnicki & Elderfield, 1993; James et al., 1995), and much of the Fe2+ that escapes sulphide precipitation is rapidly and spontaneously oxidized in well-oxygenated seawater making it difficult to evaluate the bacterial contribution to the redox transformations of hydrothermal Fe in plumes (Lilley et al., 1995; Winn et al., 1995). Nevertheless, high Fe/Mn particles and Fe-encrusted capsule forms have been observed in plumes at Axial Volcano (JDFR); the physicochemical characteristics of the capsules may be responsible for the passive or surface-enhanced deposition of iron (Cowen et al., 1999). Thiosulphate, the primary product of sulphide autooxidation, may also serve as a useful energy source, but this is yet to be documented in hydrothermal plumes (Winn etal., 1995; Cowen & German, 2002). Systematic changes in biogeochemical activity are expected as discharge flow characteristics and chemical composition of the venting solution change over time (Cowen et al., 1998). Plume bacteria often show down-current maxima for specific activities. For example, McLaughlin-West et al. (1999) found the highest rates of H2 oxidation at a site 16 km away from the caldera of Axial volcano. Similarly, the maximum number of methane oxidisers was found at 2–5 km from vent sources at Endeavour Segment (de Angelis et al., 1993). In some cases, Mn oxidation rates and relative abundance of some metal depositing capsuled bacteria in chronic plumes increase with distance from vents (Cowen et al., 1999). However, the overall abundance of bacteria tends to be highest at or near the main (chronic) vent source (Cowen et al., 1999). The differing spatial distribution of Fe and Mn in some surface sediments near vents has been found to correlate qualitatively with the activity of microorganisms, which likely attests to the differential influence of these microorganisms on Fe- and Mn-mineral precipitation (Sudarikov et al., 1995). 8.2.5

Biogeochemistry of off-axis vents and seafloor basalt

8.2.5.1 Off-axis vents Little microbiological and biogeochemical work has been done on off-axis vents to date. At off-axis seamounts, development of microbial mats and

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microbial precipitation of Fe-oxide and silica deposits seem to be ubiquitous (e.g. Alt, 1988; Boyd & Scott, 1991, 2001; Bogdanov et al., 1997; Emerson & Moyer, 1997, 2002). Diffuse vents at the Lost City site support dense microbial communities and biofilm development on mineral surfaces within the carbonate structures, despite the rarity of macrofaunal assemblages (Kelley et al., 2001). Enrichment culturing in aerobic and anaerobic media yielded thermophiles and mesophiles, while preliminary DNA extractions and analyses indicated that Archaeal and Eubacterial lineages are both present at Lost City. Although there is abundant evidence of chemosynthesis at Lost City (D.S. Kelley, personal communication), the nature and the biogeochemical effects of the microbial communities present there remain to be elucidated. 8.2.5.2 Seafloor basalts The injection of fresh basalt at mid-ocean ridges also creates considerable potential for lithosphere–ocean exchanges, though not all related to hydrothermal circulation. For example, the transformation of basaltic glass to secondary products by the action of circulating fluids is one of the major mechanisms for element exchange between seawater and oceanic lithosphere (Thorseth et al., 2001 and references therein). Several lines of evidence, including the presence of DNA, elevated values of C, N and P, and alteration textures resembling microbial cells in basaltic glass have led to the hypothesis that microorganisms inhabit the oceanic crust and that they may be involved in this transformation of basalt (e.g. Thorseth et al., 1995, 2001; Fisk et al., 1998; Furnes et al., 2001). Analysis of 16S-rRNA gene sequences showed that the major microbial populations in ocean crust appear to be characteristic populations indigenous to the rock and not contaminants from seawater or sediments, and that the closest relatives of these microorganisms are from cold marine waters (Arctic and Antarctic), deep-sea sediments and hydrothermal environments (Thorseth et al., 2001). Preliminary results suggest that iron- and manganese-oxidising bacteria (e.g. Gallionella) are indeed present and active in basaltic rocks (along with Fe-Mn reducing bacteria, methanotrophic bacteria and methanogenic Archaea) as evidenced by enrichment cultures, chemical analyses of metabolic processes, morphology (via SEM) and molecular phylogeny (Pedersen et al., 2001). The driving force for microbial dissolution of glass may be a need for Fe2+ (or other reduced compounds) as a source of energy (Thorseth et al., 2001). Other nutrients in basaltic glass including phosphorus and various metals could also be beneficial to the microbes (Fisk et al., 1998). Hydrogen oxidation has also been proposed as a widespread metabolism used by subsurface microbes (Thorseth et al., 1995). However, it is not clear how the dissolution of basaltic glass might be related to this metabolism. The weathering of the basalts could also be an indirect consequence of the production of acids generated by metabolism of organic carbon, hydrogen or hydrogen sulphide in fluids circulating in the oceanic crust (Fisk et al., 1998).

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Deep-sea cold seeps Distribution, occurrences and general characteristics

Cold seeps are areas on the seafloor where chemically enriched, oxygen-depleted fluids are released into the ocean by diffusion through sediments (Aharon, 2000). The causes of the fluid discharge at seeps are varied and include geological phenomena such as tectonically induced high pore-fluid pressures, petroleum or natural gas escape, artesian flow or catastrophic erosion, and submarine slides (Aharon, 2000; Tunnicliffe et al., 2003). Over two dozen deep cold seep sites have been discovered, but only about a half of them are well known biologically (Sibuet & Olu, 1998; Tunnicliffe et al., 2003). Seeps have been documented at several tectonically active convergent margins (e.g. northeast Pacific, Japan Trench, Aleutian margin, Barbados, Peru, Mediterranean) as well as passive continental margins (e.g. North Sea, Gulf of Mexico, Arctic Basin), typically not only at bathyal to abyssal depths, but also in shallow intertidal environments (see references in Aharon et al., 1997, Aharon, 2000, Herzig & Hannington, 2000; Tunnicliffe et al., 2003). Subduction zone seeps occur both on well-developed accretionary prisms and along erosive margins in waters as deep as 6000 m (Tunnicliffe et al., 2003), though seep-like clams with chemosynthetic symbionts have been found at a depth of nearly 7400 m in the Japan Trench (Fujikura et al., 1999). Recently, methane venting and a cold-seep biological community (tubeworms and vesicomyid clams) were discovered on exposed basement rock within an area dominated by transform tectonics along the Mendocino Fracture Zone (Stakes et al., 2002). The largest number of known seep sites are found at active continental margins where the oceanic plates are being subducted beneath continents. Here, the subducting plate scrapes its sediment burden against the overlying continental plate to form a compressed wedge of sediments or accretionary prism, over the subduction zone. In these sediments, pore fluids can occupy as much as 50–70% total sediment volume, and the fluids are squeezed from the sediment through diffuse flow or by focused flow along major fault structures in the overlying sediments (Herzig & Hannington, 2000). In some cases, mud volcanoes are formed by an influx of water from deep over-pressured zones (e.g. Barbados prism; Olu et al., 1997). On some passive margins (e.g. Gulf of Mexico), ancient salt deposits below sediments are involved in a process known as salt tectonics. Because the salt has a lower density than the overlying sediments, it tends to push upwards forming deep cracks in the sediments, through which fluids and gases can escape (Kennicutt et al., 1985; Aharon et al., 1997). Seeps are typically patchy in their distribution and their spatial continuity is often limited to a radius not exceeding a few metres (Aharon, 2000). However, Olu et al. (1996) reported a 1000 m2 continuous clam field along the Peruvian active margin, suggesting that more regular and more diffuse fluid flow is also

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possible. Seepage and dependant biological communities are typically constrained to large scale tectonic features (scarps, canyons) and smaller scale faults, fractures and joints (Olu et al., 1996). Some methane seeps can be as luxuriant as vents in terms of biomass (Olu et al., 1996), a fact that can be used to remotely locate seeps on the seafloor (Aharon, 2000). Seep sediments can also be easily distinguished from normal marine sediments by their dark gray to black colour (caused by iron sulphides), the distinctive smell of hydrogen sulphide and aromatic hydrocarbons, and the presence of degraded crude oil and/or asphaltic constituents (Aharon, 2000). The overall extent of seep communities in the deep ocean is not well known because margin environments are incompletely explored (Tunnicliffe et al., 2003). Despite this lack of detailed information, Sibuet and Olu (1998) established patterns of biodiversity, biogeography, trophic behaviour and fluid dependence of seep biological communities. At seeps, the reducing substances needed for chemosynthesis are derived not from the reaction of heated crustal seawater with basalt but from the degradation of sedimentary organic matter. As methane-rich interstitial fluids migrate along faults and fractures and approach the seafloor, they can mix with nearsurface pore waters that are slightly more oxidising and contain dissolved sulphate. In these mixing zones, microbially catalysed oxidation of methane using sulphate as an oxidant can then add hydrogen sulphide to seep fluids. At the seafloor, both CH4 and H2S provide energy for microbial chemosynthesis, either by free-living and attached microbes like the giant mat-forming Beggiatoa (a filamentous sulphur-oxidizer), or by symbionts within the tissues of specialised benthic fauna (Aharon, 2000; Nelson & Fisher, 1995). Methane is likely to be the primary source of energy in high-productivity communities, but production of H2S by SRBs in the sediments also has a major role (Sibuet & Olu, 1998). The organic carbon sources at seeps (methane, petroleum, other hydrocarbon gases, solid gas hydrates) are derived primarily from accumulated sedimentary organic carbon and thus are photosynthetic in origin (Tunnicliffe et al., 2003). In addition, microbial oxidation of reduced compounds at/or above the seafloor requires photosynthetically produced oxygen. Thus, while these ecosystems are primarily based on chemosynthesis (Aharon, 2000), they are not completely independent from photosynthesis-dominated pelagic ecosystems. 8.3.1.1 Gas hydrates A common feature associated with cold seeps is the presence of gas hydrates, which are naturally occurring solids comprised essentially of natural gas, mainly methane, trapped in frozen, crystalline water (see reviews by Kvenvolden, 1993 and Buffet, 2000). The occurrence of gas hydrates is controlled by an interrelation among temperature, pressure and composition, and they are stable in solid form only in a narrow range of these conditions (Kvenvolden, 1993). Because of these restrictions, gas hydrates are common mainly in polar and deep

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oceanic cold seep environments, such as those common on continental margins (depths of >300 m; Kvenvolden, 1993). The source of methane that forms gas hydrates is ultimately buried organic matter that is generally degraded microbially, though in some cases thermogenic degradation of organic matter can be important (Kvenvolden, 1993; Herzig & Hannington, 2000). Dissolution of these metastable gas hydrates is a natural consequence of tectonic uplift of accretionary prisms at plate margins and is likely to be partly responsible for the extensive methane plumes above these sites (Herzig & Hannington, 2000). The amount of methane that is slowly released from these hydrates can be enough to support a dependent biological community (Suess et al., 1999). 8.3.1.2 Geochemical fluxes Because of the low temperatures at cold seeps and the lack of large-scale subsurface convection currents, metals are not significantly mobilised from the crust (nor sediments) by the pore fluids as they are in hydrothermal vents (Herzig & Hannington, 2000). Nevertheless, seep fluids contain abundant CH4, CO2, N2, H2S and oil, though exact compositions are as varied as the geologic settings in which seeps are formed (Tunnicliffe et al., 2003). Methane is ubiquitous in all seep settings and methane plumes are often found at and above cold seeps (e.g. Suess et al., 1998); concentrations can exceed those in mid-ocean ridge hydrothermal fluids (e.g. up to 780 μM in shallow seeps off the coast of Denmark; Dando et al., 1994b) by more than an order of magnitude (Herzig & Hannington, 2000). In fact, elevated methane concentrations can be used as a tracer to detect cold seep venting remotely (Herzig & Hannington, 2000). Authigenic minerals are commonly formed in association with seep venting. Typical precipitates include primarily carbonates (e.g. calcite, aragonite, dolomite), but also barite, which in some cases can form large deposits, such as in the Gulf of Mexico and the Peru subduction zone (Dia et al., 1993; Torres et al., 1996; Suess et al., 1998). The formation of carbonates results from excess carbonate generated by microbial processes, such as anaerobic oxidation of methane (see Section 3.3). Barite forms as the result of mixing of Ba-rich fluids upwelling from the sediment and mixing with SO42−-rich bottom water at the cold seep sites (e.g. Torres et al., 1996). The source of Ba in these deposits may be from enrichment in venting fluids, such as at the Peru convergent margin (Dia et al., 1993). However, the Ba may also be derived from previously formed, non-detrital (possibly biogenic) barite that originates in high primary productivity oceanic areas and is buried in sediments (Torres et al., 1996). Sulphate depletion in the sediment column eventually causes dissolution of the barite and remobilisation of the barium, which is transported to the venting site, where it precipitates as barite upon mixing with bottom seawater. The magnitude of fluid flow rates and the modulation of fluid flow rates over time have long been two of the most significant questions in seep research, as these are crucial variables in estimating mass flux from cold seeps (Suess

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et al., 1998). These have been difficult to measure because the flow associated with seeps can be diffuse over large areas (e.g. Olu et al., 1996), but the relative importance of such diffuse flow versus more focused flow is not known (Suess et al., 1998). Fluid flow rates determined from simple advection–diffusion modelling of temperature and chemical profiles have proven to be unreliable if applied to sites which are densely populated with bottom macrofauna. This is because some seep bivalves (e.g. Calyptogena and Solemya clams) can pump enormous amounts of water, and they can dominate fluid flow and solute transport at seep sites (Wallmann et al., 1997). Fluid flow rates in a number of convergent settings have been estimated by a variety of geophysical and biogeochemical methods, with no clear consensus emerging (Suess et al., 1998). Wallmann etal. (1997) calculated a mean value for fluid flow of 5.5L m−2 d−1 based on a biogeochemical approach using oxygen flux and vent fluid analyses, whereas Suess et al. (1998) calculated an average of 20 L m−2 d−2 based on geophysical estimates. Conversely, fluid flow rates measured directly at actively venting sites are usually skewed towards higher values, as sites with the most flow may be favoured. For example, estimates of direct fluid flow rates range between 240 L m−2 d−2 for the Aleutian Trench and 500–1700 L m−2 d−1 for the Oregon and Peru margins (Suess etal., 1998). In addition, sea level differences, on a tidal timescale, can vary seepage rates by an estimated 4–7% (Boles et al., 2001). This wide range in estimates of fluid flow at cold seeps points to our general lack of detailed knowledge about fluid flow at seeps. 8.3.2

Biogeochemistry of seep sediment pore fluids

Although microbial processes are widely accepted as important components of cold seep activity, few studies have attempted to quantitatively establish the significance of hydrocarbons (including CH4) in fuelling microbial productivity (see review by Aharon, 2000). Similarily, few attempts have been made so far to identify and quantify these microbes either directly through culture studies, or indirectly by radioisotope assays or molecular methods (Aharon, 2000). Nevertheless, significant fluxes of various C, S and N compounds dissolved in seep pore fluids likely reflect various microbial processes (Table 8.2). Aerobic and anaerobic methane oxidation, and perhaps sulphate reduction (coupled to oxidation of methane or inorganic substances), are likely the most important chemosynthetic processes at seeps where methane is initially the most abundant reducing substance in the fluids (Aharon & Fu, 2000; Tunnicliffe et al., 2003). However, the mechanisms and rates of carbon (including methane hydrates) transformations, transport, and burial, and exchange with the open ocean at ocean margin seeps remain to be determined. The most compelling evidence for microbial utilisation of seep hydrocarbons as energy and nutritional carbon sources, and of the pathway of carbon flow through the seep fauna and carbonates, comes from the distribution of stable carbon isotopes at hydrocarbon seeps in the Gulf of Mexico (Aharon, 2000).

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Seep hydrocarbons are significantly depleted in 13C compared to other marine carbon sources and remain practically unchanged during C cycling (Aharon, 2000). Thus, tube worms hosting aerobic sulphur-oxidising bacteria and the chemolithotrophic sulphur oxidiser Beggiatoa sp. are clearly distinguishable as a group from the methanotrophic symbiont-hosting mussels. The latter have relatively low δ13C values that overlap with those of methane, thus confirming their use of 13C-depleted methane as a carbon source (Aharon, 2000). The heavier δ13C values of the sulphur-based chemosynthetic group are consistent with the use of carbon derived from either crude oil fractions, the residual pool of heavy (fractionated) methane and/or straight-chain, aliphatic fractions (C2–C5 gases derived from sublimating gas hydrates; Aharon, 2000). However, this incorporation of heavy carbon into biomass is not necessarily observed at seep environments outside of the Gulf of Mexico. Seeps are often characterised by the presence of isotopically light methane-derived authigenic carbonates, though the wide range of δ13C values in seep-related carbonates suggest that there is a variety of hydrocarbon-derived, inorganic carbon dissolved in the pore fluids and that this carbon is likely mixed with seawater DIC (Aharon, 2000). 8.3.2.1 Methanogenesis Obligatory anaerobic methane producers (methanogens) are believed to be ubiquitous in seep habitats, although no direct assays have been reported (Aharon, 2000). They have a restricted range of carbon and energy sources (e.g., hydrogen, acetate) and frequently form consortia with other anaerobic bacteria that can supply them with required metabolites (see below). The consortia convert organic C available in seeps into metabolites desirable to methanogens (acetate, CO2, H2) as follows: 3CH2O + 3H2O → H3COOH + 2CO2 + 4H2 CO2 + 4H2 → CH4 + 2H2O In marine sediments, the methanogenic zone is typically found below the sulphate-reduction zone because methanogens are outcompeted by sulphatereducers for nutrients (e.g. Ehrlich, 1990). The location of methanogenic activity supplying methane to seep sediments is not yet well determined, though it may vary amongst the different types of cold seeps. The biogenic origin of much of the methane in seep fluids is recognised by the large C-isotope partitioning between CH4 and CO2 of up to 70‰, with methane showing δ13C values typically between −45 and −90‰ PDB (Aharon, 2000). 8.3.2.2 Anaerobic sulphate reduction Typically, seep pore fluids are depleted in sulphate and enriched in reduced sulphur compounds, primarily H2S (e.g. Masuzawa et al., 1992; Aharon, 2000).

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This has been attributed to the anaerobic respiration by microorganisms like Desulfovibrio in seep sediments (Aharon, 2000); they use the abundant reduced carbon forms as electron donors and seawater-derived SO42− as an electron acceptor. In addition to H2S, this metabolism can also produce carbonate species and ammonia whose concentration and type depend on the nature of the reduced carbon substrates and on the buffering capacity of the environment. Sulphate and H2S often show a linear, inverse relationship in profiles of seep sediment pore fluids, further indicating the link between sulphate reduction and H2S production (Aharon, 2000). Aharon and Fu (2000) found significantly elevated rates of sulphate reduction in crude oil and methane seeps overlain by chemosynthetic fauna of bacterial mats and mussel beds, compared to non-seep sediments. The intensity of sulphate reduction can be gauged from the observed rapid decrease of sulphate below the sediment/water interface and the highly elevated values of H2S in pore fluids (Aharon, 2000). 8.3.2.3 Aerobic microbial oxidation of sulphide and methane Sulphur-oxidising and methane-oxidising bacteria (thiotrophs and methanotrophs, respectively) are important recyclers of metabolic wastes generated by anaerobes in most marine sedimentary environments (Aharon, 2000). At cold seeps, and particularly evident in association with gas hydrates, the abundance of reduced carbon compounds fuels an elevated activity of sulphate-reducing bacteria. This, in turn, leads to intense production of H2S, which fuels an extraordinary abundance of chemoautotrophs and symbionts (e.g. Nelson & Fisher, 1995; Aharon, 2000; Boetius et al., 2000; Zhang et al., 2002). The giant, filamentous bacteria Beggiatoa occasionally form mats at the sediment/water interface around seeps (e.g. Nelson et al., 1989; Aharon, 2000). They fix CO2 by oxidising H2S, with O2 as a terminal electron acceptor, e.g. 2H2S + O2 → 2S° + 2H2O In doing so, they deposit sulphur granules in the cells external to the cytoplasmic membrane. They can also use these sulphur granules as a storage product, oxidising them further (to sulphate) when sulphide is limiting (Ehrlich, 1990). Also common at seeps are the S-oxidising symbionts living in vestimentiferan tube worms and various clams (Aharon, 2000; Tunnicliffe et al., 2003). These symbionts generally fix CO2 by the Calvin-Benson cycle and share some of the carbon they assimilate with their host. Methanotrophs oxidise methane, with O2 as the terminal electron acceptor, e.g. CH4 + 2O2 → CO2 + 2H2O

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They live either as symbionts in mussels or as free-living aggregates at seeps (see references in Aharon, 2000). Although methane is their main metabolite, some methanotrophs may also assimilate up to 30% of their carbon as CO2 (Ehrlich, 1990). 8.3.2.4 Anaerobic oxidation of methane Although the oxidation of methane in anoxic sediments has long been suspected based on geochemical evidence (e.g. Hoehler et al., 1994), no organisms that can consume methane anaerobically have yet been isolated and little is known about the physiology, biochemistry and identity of such organisms, since all well-known methanotrophic (methane consuming) microbes are aerobic bacteria (Boetius et al., 2000; Orphan et al., 2001, 2002). However, recent geochemical, isotopic and microbiological data indicate that a coupling exists between sulphate reduction and anaerobic methane oxidation (Hinrichs et al., 1999; Aharon & Fu, 2000; Teske et al., 2002) and that novel Archaea and sulphatereducing bacteria are involved (Boetius et al., 2000; Hinrichs et al., 2000; Orphan et al., 2001, 2002; Teske et al., 2002). It has been proposed that anaerobic oxidation of methane involves at least two types of microorganisms operating together: (1) methanogens, running in reverse, oxidise methane to CO2 and H2, or to acetate, and (2) SRBs reduce sulphate by oxidising H2 or acetate from methane oxidation (e.g. Hoehler et al., 1994). Recent work on cell-specific stable C isotope analyses and molecular phylogeny has confirmed that this process is mediated by several organisms, including novel Archaeal methanogens and sulphatereducers (using an unknown intermediate substrate). These microbes have been found individually and as consortia with the methanogens in the core surrounded by SRBs (Boetius et al., 2000; Orphan et al., 2001, 2002; Fig. 8.7). The methane-consuming Archaea are phylogenetically distinct from known methanogens and they are highly diverse (Orphan et al., 2002; Teske et al., 2002). Such microbial consortia are also abundant in gas-hydrate-rich sediments at Hydrate Ridge, where sulphate reduction rates were the highest ever measured in cold marine sediments (140 mmol m−2 d−1). Boetius et al. (2000) concluded that, at Hydrate Ridge, sulphate reduction is fuelled by high methane fluxes from below, while organic deposition from surface waters is not a significant substrate source for SRBs. Similar findings are reported from the Gulf of Mexico (Aharon & Fu, 2000) and Eel River Basin (Orphan et al., 2001). The overall syntrophic process involves a transfer of electrons from methane to sulphate, e.g. CH4 + SO42− → HCO3− + HS− + H2O

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Fig. 8.7 Fluorescence in situ Hybridisation (FISH) overlay image of microbial aggregate involved in anaerobic methane oxidation. Different microbes were stained using different oligonucleotide probes. Core is ANNM-2 Archaea, surrounded by sulphate-reducing Desulfosarcina imaged by laser scanning confocal microscopy. From V.J. Orphan et al. in Proc. Natl. Acad. Sci. USA, Vol. 99, 7663–7668, 2002. Reproduced with permission. Copyright National Academy of Sciences (2002).

Methane-derived intermediates (possibly acetate or CO2) and hydrogen from methane-consuming Archaea are transferred to their sulphate-reducing bacterial partners (Orphan et al., 2001). Teske et al. (2002) recently demonstrated that anaerobic methane oxidation occurs in Guaymas Basin. Although this site is influenced by active hydrothermal venting, the abundant hydrocarbon-rich sediments make the system similar to cold seeps. These authors suggest that such anaerobic methanotrophy is likely a significant and widely occurring process at methane-rich hydrothermal vents. 8.3.3

Microbial carbonates

Massive carbonates and carbonate-cemented sandstone deposits (e.g. Fig. 8.8) are common features at cold seeps, primarily because of high microbial activity that can decrease CO2 from pore fluids and increase alkalinity. Hydrocarbonconsuming bacteria (e.g. methane oxidisers) produce unusually high levels of alkalinity and DIC in the pore fluids which drive the carbonate system to supersaturation and consequently to carbonate deposition (Wallmann et al., 1997; Clari & Martire, 2000; Aharon, 2000). Sulphate reduction coupled with methane oxidation also affects alkalinity and favours carbonate precipitation (e.g. Aharon, 2000; Zhang et al., 2002). Development of high alkalinity in seep pore fluids resulting from release of OH− and HS− during microbial sulphate reduction would also enhance carbonate precipitation because of higher pH

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Fig. 8.8 Ring-shaped carbonate concretion from a cold seep. The unusual shape of the concretion is likely a result of pressure and sediment outgassing effects. Reproduced with permission. Mark Leet for MBARI © 1997.

shifting the carbonate equilibria toward the carbonate ion (Aharon, 2000). Carbon dioxide fixation (removal of CO2) by chemolithotrophic microbes using the Calvin-Benson cycle can also drive carbonate precipitation from a pore fluid enriched in bicarbonate (Aharon, 2000), e.g. 2HCO3− → CO2 + H2O + CO32− Botryoidal aragonite cements are among the most common textural features in seep carbonates (e.g. Clari & Martire, 2000), and a microbial origin is inferred by anomalously negative δ13C values, suggesting a methane-derived C-source; presence of dead bacterial cells filling dark centres of botryoids as seedling sites and coexistence with opaque sulphides suggest that anaerobic sulphate reduction also occurred in pore fluids, along with carbonate ion enrichment (see references in Aharon, 2000). Typically, these carbonates also show laminated build-up structures similar to stromatolites. Microbial processes could also lead to episodic carbonate dissolution as CO2 produced during methanogenesis and methane oxidation could lower the

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degree of saturation of pore fluids with respect to carbonates by decreasing pH (Aharon, 2000). These ubiquitous features of seeps preserve detailed records of microbial processes long after seepage has ceased and the chemosynthetic communities have vanished (e.g. Clari & Martire, 2000). Unusual carbonate deposits in the rock record (as far back as at least early Cretaceous) have been recognised as having precipitated around cold seeps of methane-rich fluids (Beauchamp et al., 1989; Aharon, 1994; Clari & Martire, 2000). The carbonate minerals in these deposits are strongly depleted in 13C (−25 to −40‰, or lower), which is interpreted to be a result of the incorporation of isotopically light CO2 from bacterial degradation/oxidation of methane. Additional evidence of bacterial activity and mediation of carbonate precipitation in ancient deposits include the presence of abundant pyrite in carbonate groundmass suggesting activity of sulphatereducing bacteria; unusual structures and fabrics not commonly seen in marine siliciclastic rocks, carbonate cements and internal sediments; and a peloidal fabric and dumbell-shaped dolomite spheroids similar to those linked to microbial activity (Clari & Martire, 2000).

8.4

Stability and perturbations of seafloor hydrothermal vent and cold seep systems

Seafloor hydrothermal and cold seep systems host biological communities that are extremely adapted to their environment and that respond rapidly to environmental changes and perturbations. Thus, these systems can be used as models for understanding various biogeochemical and ecological processes, and for testing models and hypotheses. Also, these systems are important components in global biogeochemical cycles. Thus, understanding the stability of these systems and their response to natural or anthropogenic perturbations (i.e. resilience) is essential. 8.4.1

Geological stability of vents and seeps

Individual deep-sea hydrothermal vents are generally short-lived. However, the heat from the cooling of deep oceanic crust and mantle can support hydrothermal systems episodically for tens of thousands of years (Kelley et al., 2002). Long-lived, stable hydrothermal systems occur in young crust on the axis, in older crust away from the spreading axis and in cooler ridge flanks in crust up to ~65 Ma (see references in Kelley et al., 2002). The TAG site on the MAR may have seen hydrothermal activity as far back as 150 000 years (Lalou et al., 1995). Older systems tend to support non-vent deep-sea communities which contain numerous filter feeders (e.g. Explorer Ridge; Embley et al., 2002; Léveillé et al., 2002), and extinct Fe-sulphide-rich chimneys may support

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Fe-S-oxidising bacteria over long periods of time, even after hydrothermal activity has ceased. Cold fluid seepage at any given locality can have a geologically short life span. However, several biological observations (e.g. trophic complexity and high species richness) in modern seep systems hint at sustained seepage at several sites (e.g. Olu et al., 1996). Aharon et al. (1997) proposed that the seepage in the northern parts of the Gulf of Mexico has occurred from at least the late Pleistocene to the present; ages (based on radiometric dating) of carbonate rocks and shells from extinct shallow sites there indicate that they are as old as 200000 years. However, it is not clear how continuous seepage might have been over this interval. Once seepage stops, so does the chemosynthetic activity of organisms. The only vestige is the long-lasting carbonate buildups which contain a detailed record of seep history. 8.4.2

Future perturbations related to resource extraction

8.4.2.1 Hydrothermal sulphides Large seafloor deposits of polymetallic massive sulphides formed by hydrothermal venting are considered as analogues of some terrestrial deposits, and the metal content especially that of Cu, Zn, Pb, Ag, Au, and Ba in these deposits can be high (Scott, 1997). Mining companies have taken an interest in these deposits, with one such company holding exploration licences offshore of Papua New Guinea (Scott, 2001). Underwater mining is not new (e.g. diamonds off the coast of southern Africa) and offshore oil exploitation is now widespread. Thus, it appears likely that mining of seafloor minerals will progress in the near future, despite the scientific and technological challenges that need to be overcome. The main hurdle at present is the mediocre economic climate for base metals. Large scale mining of seafloor polymetallic sulphide deposits could significantly disturb all aspects of the hydrothermal system including fluid flow, mineralisation, plume composition and macro/microbiological activity. Habitat would certainly be modified significantly and some organisms would definitely be killed directly by the mining activity. However, mining would only be localised in very limited areas where deposits of commercial size are known to occur (Scott, 2001). Vent organisms are regularly disturbed by natural perturbations caused by volcanic eruptions, earthquakes and chimney growth. These communities have shown a remarkable ability to recover from such perturbations. However, while it may be tempting to apply this resilience argument to considerations of mining impact, it is important to point out that the mother populations that permit repopulation after perturbation are themselves particularly vulnerable to mining. Long-lived mother populations may be critical to the maintenance of vent species’ biodiversity within a region, especially at larger, longer-lived

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hydrothermal sites where diversity tends to be greatest. These long-lived hydrothermal sites are also the most likely locations for accumulation of large sulphide deposits and therefore will be prime targets for mining (Scott, 2001). As well, many localised species may not have a nearby mother population, or they may be unable to recolonise the altered substratum after mining. In the latter case, only the establishment of protected areas would prevent eradication of species (e.g. Dando & Juniper, 2001). 8.4.2.2 Cold seeps Gas hydrates, such as those commonly found at cold seeps, are also of economic interest, for three main reasons: (1) they represent huge reserves of methane, (2) methane is a relatively clean fossil fuel energy source with a high energy content compared to oil and coal, and (3) they are distributed worldwide, though knowledge of their occurrence is incomplete (Kvenvolden, 1993). Theoretically, one cubic metre of CH4 hydrate can contain up to 164 m3 of methane gas at standard temperature and pressure conditions (Kvenvolden, 1993). Gas hydrates may also serve as traps for gaseous methane, thus creating even larger reserves (Kvenvolden, 1993). The amount of methane in marine gas hydrates worldwide is not well constrained. Estimates range from ~1015 to 1017 m3 of methane (Kvenvolden, 1988; Gornitz & Fung, 1994; Kvenvolden, 1999). Even at the lower end of the range, reserves of methane below the seafloor would exceed, by a factor of ~2, all presently known fossil fuel deposits (Kvenvolden, 1993). Given these facts, gas hydrates could serve as an important future energy resource (e.g. Kvenvolden, 1999). However, little is known about the possible effects of mining seafloor methane on cold seeps. Any disturbance could potentially cause a release of pressure, which could then trigger a massive release of methane through destabilisation of the hydrates. This could have important climate implications because of methane’s strong global warming potential, which is estimated to be 23 times that of CO2 over a 100-year time horizon (IPCC, 2001). 8.4.3

Response of cold seeps and gas hydrates to global warming

Because huge quantities of methane are buried in the seafloor, mainly as gas hydrates, there are concerns about potential climate effects, especially because methane has a greater greenhouse effect than CO2. The concentration of atmospheric methane is also presently increasing at a more rapid rate than that of CO2 (IPCC, 2001). Therefore, the abundance, distribution (in space and time) and the stability of methane hydrates have important implications for future and past global climate changes. Global warming may lead to instability of gas hydrates, as a result of changes in pressure and temperature associated with global sea level rise and warming; thus, a positive climate feedback mechanism may result (e.g. Kvenvolden,

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1988; Kvenvolden, 1999). In this scenario, the warmer climate would cause hydrates to release methane gas, which would eventually make its way to the atmosphere, where it would contribute to even greater warming. It has been suggested that such a mechanism has acted in the past to rapidly change earth’s climate (e.g. Nisbet, 1990). For example, Dickens et al. (1997) suggest that the warm Paleocene may represent a past analogue for possible future effects on global biogeochemical cycles. It should be noted that such scenarios are highly speculative because it is not even known how gas hydrates behave in the present climate regime (Kvenvolden, 1999). The amount of methane likely to be released as a result of future global warming is not known. One point to consider is that the coupled activities of methane-oxidising and sulphate-reducing organisms contribute to the sequestration of carbon as buried organic carbon and authigenic carbonates. Thus, methane released into the water column may be oxidised before it ever reaches the atmosphere (e.g. Kvenvolden, 1999). Microbial consortia of Archaea and SRBs may consume as much as 80% of the methane produced in the oceans each year (Orphan et al., 2001), thus avoiding huge fluxes of methane to the oceans and ultimately to the atmosphere (Zhang et al., 2002). Thus, this microbial activity may represent a biogeochemical damper preventing a runaway greenhouse effect. However, this mechanism is yet to be considered in studies that attribute past global climate warming to massive release of methane from gas hydrates (Kvenvolden, 1999).

8.5

Future work

Considering that deep-sea environments are some of the least explored parts of our planet, we have gained a wealth of information about them in only a short time span. Research on deep-sea hydrothermal vents, and to a lesser extent on cold seeps, has matured over the last decade or so, as evidenced by the large number of extensive review papers in many related disciplines. Despite these advancements, many unanswered questions remain. There is now a strong need to move from a descriptive to a quantitative mode; such studies on biogeochemical processes have been few in number. Future prospects for research on deep-sea hydrothermal vents and cold seeps are very promising. New technical approaches and exploration methods (e.g. autonomous vehicles like ABE; Embley et al., 2002) will likely benefit this research. Molecular methods will help us to better understand the diversity and possibly the metabolisms of un-culturable environmental microorganisms. Long-term seafloor observatories like NeMO (Embley & Baker, 1999) and NEPTUNE (Delaney et al., 2000) will allow us to more easily study episodic events like volcanic eruptions at ridges and massive methane releases from seeps. As well, these observatories will provide data in the form of long-term time

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series, which are presently lacking for these systems, especially for issues pertaining to the dynamic nature of vent fluid compositions and microbial activity following dyking-eruptive events and associated heat and element fluxes to the ocean. Since hydrothermal element outputs are generally derived by multiplying the mean vent fluid composition for that element by an estimate of global hightemperature fluid flux, accurate estimates of the latter are essential. Unfortunately, such estimates are not presently well constrained. In addition to episodic events, the relative importance of axial and off-axis fluxes needs to be quantified. This could have important consequences for estimating global fluxes from hydrothermal vents (Elderfield & Schultz, 1996). Increased exploration and subsequent studies of off-axis sites (e.g. Kelley et al., 2001) will contribute greatly to this poorly studied issue. At vents, carbon dynamics and the role of heterotrophy need to be elucidated. Cowen and German (2002) advocate the need for particle flux measurements in conjunction with other physical and biogeochemical measurements, and rate experiments in hydrothermal plumes. The fate of plume organic carbon exported from vent systems plus that produced in situ needs to be characterised. Measurement of non-conservative species such as Mn, CH4 and H2 is important for understanding underlying geochemistry of the hydrothermal system and the active geochemical and microbial processes that control the transformation, dispersion and fate of these species (Cowen et al., 1998). An understanding of the geomicrobial processes active within plumes is crucial to improve the diagnostic power of plume measurements in order to address temporal and spatial variability, for improved resolution in assessing hydrothermal chemical fluxes and for greater appreciation of the impact of hydrothermal discharge on ocean chemical budgets and biological communities (Cowen and German, 2002). Numerous studies suggest that the zone within the highly permeable, porous and fractured mid-ocean ridge crust that is occupied by microorganisms may be extensive. However, many key questions remain concerning the subsurface biosphere at mid-ocean ridges, especially in relation to its extent, both on- and off-axis, the identity and nature of the microorganisms there and their levels of activity in the subsurface. Solving these problems will require direct observations and sampling of the subsurface environment and a better understanding of the overall subsurface geology and fluid flow. Although eruptive events may provide us with glimpses (or windows) of a subsurface biosphere, a more systematic and direct approach to its study is needed (Tunnicliffe et al., 2003). The interest in the thermostability of hyperthermophilic enzymes in terms of biotechnological applications (e.g. Cowan, 1995) will continue to stimulate microbiological research at vents. Fe-oxyhydroxide-silica precipitation associated with deep-sea hydrothermal vent microbes appears to be ubiquitous. It remains to be determined if this is

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simply a result of the prevailing physicochemical conditions in these environments, or if the microbes actively or passively facilitate formation of these minerals. Proving bacterial involvement in any mineralisation process is difficult. Further studies elucidating microtextures, especially via SEM, will do little to advance this area of research. While this may not be realistic for older, more mineralised deposits, it is certainly possible for young, actively forming deposits. Standard SEM analyses are most useful for studying wellmineralised samples, and are poor tools for elucidating the earliest stages of mineralisation of bacterial cell surfaces. For the latter, environmental or variablepressure SEM analysis is more suitable as these instruments can be used to examine hydrated, biological samples such as biofilms and mats, which are common on mineral surfaces in seafloor hydrothermal deposits (see reviews by Karl, 1995; Juniper & Tebo, 1995; Kelley et al., 2002). At the very least, TEM investigations (e.g. Fortin et al., 1998; Boyd & Scott, 2001) are needed to demonstrate the presence of cellular membranes or cytoplasms in such mineral deposits. While morphological evidence can be convincing and much can be learned from electron microscopy and molecular phylogeny, rigorous demonstration of bacterial Fe-oxidation requires isolation of organisms in Fe-enrichment cultures and observation of Fe accumulation or release from sulphides during in vitro growth coupled with CO2 fixation (Juniper & Tebo, 1995). Only then can the quantitative importance of microbial iron oxidation be determined. The question of microbial sulphide weathering at vents is currently being seriously addressed by a variety of combined approaches, including textural and isotopic biomarkers, water-rock interaction experiments and in situ colonisation experiments (Edwards & McCollom, 2001; Léveillé & Juniper, 2002; Léveillé et al., 2002). Sulphide weathering may prove to be a quantitatively important, long-term biogeochemical process at mid-ocean ridges. In order to assess the impact of seeps on marine biogeochemical cycles, particularly their role in the cycling of C, S and N, we need to understand better hydrocarbon fluxes, temporal and spatial frequency of seepage, rates of anaerobic and aerobic bacterial respiration and transport rates of metabolites (Aharon, 2000). At present, there are few pore fluid chemistry analyses reported in the literature – many more are needed to assess the influence of microbial processes on pore fluid chemistry (Aharon, 2000). Recent technological advancements in determining direct fluid flow rates (e.g. Boles et al., 2001; Tryon et al., 2001) appear to be promising while measurements of fluid flow rates are lacking. A link needs to be established between pore fluid chemistry alterations attributed to specific bacterial processes and unique textures in carbonate deposits containing bacteria imprints. Such an association can be best documented by multidisciplinary investigations of modern seeps which have been explored insufficiently.

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Conclusion

Research on deep-sea hydrothermal vents and cold seeps is still relatively young, both systems having been discovered within the last 25 years. Nevertheless, the importance of vents and seeps in marine biogeochemistry is now widely recognised, though still poorly understood. In the past decade or so, we have come to understand that there exist important linkages between geological, chemical and biological processes at deep-sea hydrothermal vents and cold seeps. Many of the processes associated with these systems provide energy and nutrients that support abundant and diverse microbial communities in a variety of environments. These microbes, in turn, exert a strong influence on a variety of processes from fluid flow and chemical fluxes to mineral weathering and precipitation. Microbes, we now know, can be important geological agents. References Aharon, P. (2000) Microbial processes and products fueled by hydrocarbons at submarine seeps, in Microbial Sediments (eds R.E. Riding and S.M. Awramik), Springer-Verlag, Berlin, pp. 270–281. Aharon, P. and Fu, B. (2000) Microbial sulfate reduction rates and sulfur and oxygen isotope fractionations at oil and gas seeps in deepwater Gulf of Mexico. Geochimica et Cosmochimica Acta, 64, 233–246. Aharon, P., Schwarcz, H.P. and Roberts, H.H. (1997) Radiometric dating of submarine hydrocarbon seeps in the Gulf of Mexico. Bulletin Geological Society America, 109, 568–579. Al-Hanbali, H.S., Sowerby, S.J. and Holm, N.G. (2001) Biogenicity of silicified microbes from a hydrothermal system: relevance to the search for evidence of life on earth and other planets. Earth and Planetary Science Letters, 191, 213–218. Alt, J.C. (1988) Hydrothermal oxide and nontronite deposits on seamounts in the Eastern Pacific. Marine Geology, 81, 227–239. Alt, J.C. (1995) Subseafloor processes in mid-ocean ridge hydrothermal systems, in Seafloor Hydrothermal Systems: Physical, Chemical, Biological, and Geological Interactions (eds S.E. Humphris, R.A. Zierenberg, L.S. Mullineaux and R.E. Thomson), Geophysical Monograph 91, American Geophysical Union, Washington, DC, pp. 85–114. de Angelis, M.A., Lilley, M.D., Olson, E.J. and Baross, J.A. (1993) Methane oxidation in deep-sea hydrothermal plumes of the Endeavour Segment of the Juan de Fuca Ridge. Deep-Sea Research I, 40, 1169–1186. Avery, E.R., Zierenberg, R.A. and Nelson, D.C. (2000) Seafloor oxidation of massive sulfide deposits, Ridge Theoretical Institute – The Subsurface Biosphere at Mid-Ocean Ridges, July 27–August 1, Big Sky, Montana, p. 6. Baker, E.T. (1995) Characteristics of hydrothermal discharge following a magmatic intrusion, in Hydrothermal Vents and Processes (eds L.M. Parson, C.L. Walker and D.R. Dixon), Geological Society Special Publication No. 87, pp. 65–76. Baker, E.T. (1998) Patterns of event and chronic hydrothermal venting following a magmatic intrusion: new perspectives from the 1996 Gorda Ridge eruption. Deep-Sea Research II, 45, 2599–2618. Baker, E.T. and Massoth, G.J. (1988) Characteristics of hydrothermal plumes from two vent fields on the Juan de Fuca Ridge, Northeast Pacific Ocean. Earth and Planetary Science Letters, 85, 59–83. Baker, E.T., German, C.R. and Elderfield, H. (1995) Hydrothermal plumes over spreading-centre axes: Global distributions and geological inferences, in Seafloor Hydrothermal Systems: Physical, Chemical, Biological, and Geological Interactions (eds S.E. Humphris, R.A. Zierenberg,

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9

Influence of nutrient biogeochemistry on the ecology of northwest European shelf seas Paul Tett, David Hydes and Richard Sanders

9.1

Introduction

In coastal seas, as in the oceans as a whole, the main primary producers are micro-algae and cyanobacteria, constrained to be small in order to float in the illuminated region near the sea surface and to be able to assimilate key mineral compounds – nutrients – from low concentrations in sea water. The organic matter made, and the chemical elements harnessed by, these phytoplankters, supports a pelagic food web leading by way of herbivorous protozoa and small invertebrates to carnivorous larger invertebrates, fish, and sea mammals and birds. A related benthic food web leads from microbial decomposers of sedimented organic matter by way of protozoa and invertebrates to bottom feeding fish and other carnivores. The focus of this chapter is the microplankton (the phytoplankton and its bacterial and protozoan consumers) of shelf seas, and our starting point is the observation of Redfield (Redfield, 1934, 1958; Redfield et al., 1963) that plankton tends to a standard chemical composition, which in terms of ratios of atoms is C, 106 to N, 16, to P, 1. Redfield saw that since it was mainly the mineralisation of plankton-produced organic matter that refreshed the inorganic compounds of nitrogen and phosphorus in sea-water, the marine (molar) ratio of nitrate to phosphate should also tend to 16:1. That is, the composition of plankton and these aspects of the composition of seawater were part of a cycle, each determining the other. Modern understanding brings other elements, such as silicon and iron, into these ratios and invokes a wider range of biogeochemical processes to explain, for example, typical levels of phosphorus in the sea. Nevertheless, such analyses imply either that all planktonic photosynthesisers are essentially the same, or that only a few types need to be recognised, such as diatoms for their role in silica cycling and the small flagellated algae covered in limy plates called coccoliths for the part that these plates play in carbon cycling. However, diversity is one of the most obvious features of most collections of organisms, and coastal plankton is no exception to this. One possible explanation is that such diversity is a mere biological epiphenomenon, the result of life’s tendency to vary without biogeochemical cause or consequence. In this chapter, however, we explore the extent to which microplanktonic diversity

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does have a biogeochemical explanation, and what the ecological consequences of that biogeochemical forcing might be. We consider whether deviations from the Redfield ratio (defined to include silicon) may be at least partly responsible for the differential success of different kinds of planktonic algae and cyanobacteria, and thus contribute to that part of the process of human-induced eutrophication that includes an undesirable disturbance to . . . the balance of organisms (C.E.C., 1991). We focus on temperate shelf seas because these are not only best studied, but also most subject to anthropogenic influence. Most of our examples are taken from studies in northwest European waters, because it is these that we know best. The aim of this chapter is to review recent improvements in knowledge of the sources and sinks of nutrients in the shelf seas of northwest Europe, and to consider, especially, the influence of ratios of nutrient elements on the floristic composition of the phytoplankton in these seas. We deal mainly with the macronutrient elements nitrogen, phosphorus and silicon, and to some extent with iron. The acronym DAIN (Dissolved Available Inorganic Nitrogen) is a convenient way to refer to nitrate, nitrite and ammonium, excluding di-nitrogen which is not available to most phytoplankters. 9.2 9.2.1

Nutrient cycles Macronutrient element availability

Phytoplankton can take up a wide range of nitrogen species from solution. The principle forms are nitrate, nitrite and ammonium, although some species are capable of fixing dissolved nitrogen directly into new biomass, and some forms of organic nitrogen, especially urea and dissolved free amino acids, may play important roles. Nitrogen can occur in a variety of oxidation states ranging from +5 (nitrate) to −3 (ammonium). The conversion of nitrogen through these various oxidation states is a biologically mediated process, and some of the reactions are light sensitive, especially the conversion of ammonium to nitrate (nitrification). Plants generally require reduced nitrogen and are frequently considered to prefer ammonium to nitrate or nitrite for this reason. In addition, the acquisition of ammonium is a passive process driven by a diffusive gradient into the cell. In contrast, the acquisition of nitrate is a two step enzymatically driven process involving the enzymes nitrate reductase and nitrite reductase, both of which are energetically expensive to synthesise. In addition to nitrogen, algae also have an absolute requirement for phosphorus. The principle form of phosphorus available to algae is orthophosphate. Some classes of algae, most notably diatoms, have an absolute requirement for silicate which they use in cell walls and other structures. Under the low nutrient summer conditions that follow the spring plankton bloom, the single largest pools of fixed nitrogen and phosphorus in coastal

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waters are, frequently, the dissolved organic pools. For example, a transect of stations from the inner Thames estuary to the centre of the southern North Sea shows that the nitrate pool declines proceeding offshore, whilst the dissolved organic nitrogen (DON) pool simultaneously assumes greater significance (Sanders et al., 2001). The quantitative significance of the organic species in phytoplankton nutrition is uncertain (Caron et al., 2000). Whilst it is widely established that phytoplankton can take up a wide range of nitrogeneous organic compounds from solution, the coincidence of DON domination of the fixed nitrogen pool with the lowest phytoplankton biomass suggests that the DON present at this time is relatively unimportant for phytoplankton nutrition. Possibly, there is a small pool of rapidly recycled organic matter fuelling the microbial loop superimposed on a large pool of biologically inaccessible material which may be rendered accessible via photoloysis. The transport of terrestrialderived, riverine, dissolved organic material through estuarine systems and into the coastal zone has been reported (Mantoura & Woodward, 1983; Seitzinger & Sanders, 1997; Minor et al., 2001). Dissolved organic phosphorus can be accessed via the enzyme alkaline phosphatase, and the analysis of levels of this enzyme has shown that DOP utilisation may be quantitatively significant in the coastal zone (Suzumura et al., 1998; Gambin et al., 1999). Rather few studies have directly considered the role of particulate organic matter (POM) as a substrate for micro-algal growth. Nevertheless, given that much of the recycled nutrient pool is derived from the remineralisation of algal biomass, the POM pool must represent a pool of utilisable organic matter. Isotopically based studies (Fichez et al., 1993) suggest that terrestrial POM persists into the coastal zone, hinting at a relatively inaccessible nature. The mineralisation of sedimented organic material regenerates a substantial proportion of the nutrients which replenish the North Sea ecosystem during winter; thus, in this context at least, POM is a valuable algal substrate (Nedwell et al., 1993; Hydes et al., 1999). Consideration of the particulate cycle in coastal waters is important for several reasons. Firstly, nutrient transport may become decoupled from water transport if a significant fraction resides in the particulate phase. Secondly, particles might represent a reservoir of semiavailable nutrients that erosion or resuspension could introduce to the water column where desorption or remineralisation might occur. In addition to the biologically driven assimilation of nutrient species into the particulate material discussed above, we must also consider the abiotic sorption of nutrient species to inorganic sedimentary particles, which frequently occur in great numbers in energetic coastal waters (Jickells, 1998; Prastka et al., 1998; Jickells et al., 2000). The strength of this sorption is defined by the distribution coefficient, Kd. Values of around 50 000 L kg−1 are typical for phosphate implying that the equilibrium between the dissolved and particulate phase is heavily biased

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towards the particulate phase at moderate particulate loads. Kd values for nitrate and ammonium are such that the inorganic sorption by particles is an unimportant process. The practical consequences are that, in estuarine systems that have been heavily industrialised or that drain intensively populated or farmed catchments, there is likely to be a high particulate load, as salt marshes have been removed, and also high phosphate loads. The high sediment loads impede primary production and consequently the particle–water interaction phenomenon is likely to be the dominant influence on phosphorus movement. This is indeed observed in some regions (Fichez et al., 1992; Shaw et al., 1998). Very high removal of P is seen in the low salinity reaches of many estuaries (Prastka et al., 1998). However, the fate of the P-enriched material is important, since burial will remove P from the system and particle flushing may result in desorption in more saline waters where the particulate load is lower. One consequence of industrialisation is that not only are the particulate loads higher but also the sedimentation is lower, thus flushing is the likely fate. The desorption of large amounts of P in offshore regions means that the local DAIN:DIP ratio may be a poor indicator of the potential limiting factors. 9.2.2

Sources of macronutrients

Coastal shelf seas receive nutrients from a diverse range of sources including oceans, rivers and atmosphere, local water-column and sea-bed recycling of biologically produced organic compounds. Perhaps, the region where these inputs have been best documented is the North Sea. We, therefore, initially review the various delivery mechanisms themselves, before attempting to describe their relative importance to the nutrient budget of this system. Whether or not other coastal waters have experienced increases in nutrient concentrations is a moot point – as can be seen in the case of the Irish Sea (compare: Allen et al., 1998; Gowen et al., 2002). Over the last 50 years, all watersheds with river discharges into the North Sea have been subject to increased inputs of both nitrogen and phosphorus, due to the more systematic collection of sewage and more intensive agricultural practices. This has resulted in increase of one to two orders of magnitude in the concentrations of dissolved nitrogen and phosphorus in the rivers draining these catchments and entering estuarine systems. Analyses (Nedwell et al., 1999; Jickells et al., 2000) suggest that a fraction of the material entering estuaries is permanently removed (via sedimentation or denitrification) or transformed into particulates. The impact of these processes on net estuarine nutrient fluxes (the filtering effect) is generally low in short residence time, hypernutrified estuaries. Most North Sea estuaries are of this sort, and thus a substantial fraction of the terrestrial nutrients in runoff is delivered to coastal waters in one form or another. The transfer of ocean properties into shelf waters is not straightforward. The observations discussed by Pingree et al. (1999) demonstrate the degree to

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which the physical processes at the shelf break insulate the northwest European shelf from the Atlantic Ocean. Nevertheless the intermediate salinity of North Sea waters, lying between oceanic and freshwaters, implies that the transfer of oceanic properties across the shelf break does occur. An important question is how fast this transfer occurs and in particular, the extent to which this transfer occurs over timescales less than those which remove nutrient species from the coastal shelf ecosystem. The issue of atmospheric deposition as a transport mechanism for nutrients to coastal shelf seas is a relatively recent one. However, it has now emerged as being of great importance for the following reasons. Firstly, this method is very significant in a quantitative sense (Rendell et al., 1993; Asman et al., 1995). Secondly, the deposition of material frequently occurs some way out to sea (Jickells, 1995), where the impact of biological processes in the immediate coastal zone has already reduced nutrient concentrations to undetectable levels. Particular sources of nutrients include agricultural emissions from livestock and motor vehicle exhaust fumes. The flux of nutrients to coastal seas via groundwater discharge is a subject we know almost nothing about. In the last ten years, work has been done in the North Sea by a number of national and international projects. A key project was the UK Natural Environment Research Council’s North Sea project (NSP). A unique feature of the NSP was that a consistent grid of stations was worked in the southern North Sea once a month for fifteen months (Howarth et al., 1993). From this, the winter standing stock of nitrate and silicate was calculated. The inputs to the system over an annual timescale were then calculated to determine whether they can account for this standing stock. The river input was calculated for the appropriate periods for the 11 main river systems flowing into the southern North Sea. This was done on the basis of a data set compiled at SOC from data at BODC, Rijkswaterstaat and ARGE. Atmospheric inputs were calculated using NSP data for wet and dry deposition (Rendell et al., 1993). Fluxes of dissolved nutrients across the sediment–water interface were estimated using the data from the six NSP sites. Net exchanges across the northern and southern boundaries (the Dover Straits and a line at 56° N) resulting from the advection and dispersion of nutrients were calculated using two model-based methods, and by simple multiplication of water flow estimates and nutrient concentrations at the boundary (Laane et al., 1993). Table 9.1 shows the resulting nutrient budget for the North Sea. In this budget, both estimates for dissolved inorganic nitrogen (DIN) and silicon fluxes are included. The silicon budget is included as a control. The processes involved in the silicon cycle are fewer than that of nitrogen, and so it might be expected that the budget can be closed more easily. The silicon cycle comprises essentially only river and ocean inputs, removal of dissolved silicon

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Table 9.1 Nitrogen and phosphorus budgets for southern North Sea

Observed load in January Change Nov to Jan (=A) River input Atmospheric Cross boundary1 Sediment exchange Total input (=B) Internal input (=A−B)

DAIN (ktonnes) N

Silicon (ktonnes) Si

960 526 82 38 12 11 143 383

990 242 54 – 4 273 331 −89

1

Cross boundary calculation by particle tracking method. Comparison of quantifiable nutrient input tonnages from different sources with the measured change in southern North Sea between November and January (Hydes et al., 1999).

into diatom cell walls in the spring bloom and then dissolution of silicon from these cell walls on the sea bed later in the year. The nitrogen budget in Table 9.1 is out of balance by about 350 kilotonnes of DIN. The question arises as to the source of the missing nitrate-nitrogen. If this is not the sediments, rivers, atmosphere or outer boundaries, then there must be a source within the water column source. There are two possible sources: dissolved organic nitrogen compounds (DON) or organic detritus suspended in the water column. DON measurements were not made during the NSP cruises and there is insufficient data available from other sources to allow a budget to be made. Consideration of the size of the DON and PON pools suggests that mineralisation of this reservoir could be large enough to supply the missing nitrate. This discussion highlights the need for more measurements of DON and suspended particulate N. 9.2.3

Sinks of macronutrients

In the short term, the most significant temporary sink for dissolved nutrients is uptake by phytoplankton. This is followed by remineralisation, which returns nutrients to solution or loss from the system by burial in sediments or export to ocean waters. These processes effectively separate N, P and Si because remineralisation returns them to solution at different rates (Officer & Ryther, 1980). In particular, Si dissolves slowly from shell material, and denitrification leads to the loss of biologically available nitrogen from the system as nitrogen gas. In temperate waters, nutrients undergo seasonal cycles with algal uptake dominating during spring and summer. Joint and Pomroy (1993) estimated annual levels of production in the different North Sea Task Force (NSTF) areas of the southern North Sea. Hydes et al. (1999) compared estimates of production with estimates of available nitrogen based on observed loads before

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the spring bloom, and inputs from rivers and the atmosphere during the productive part of the year. Levels of primary production throughout the North Sea were higher than the estimated loads of available nitrogen, suggesting that production is maintained by a recycling of nutrients within the system. The ratio of production to available nitrogen varied from two in English coastal waters to five in the German Bight. The lower production off the English coast is consistent with the higher suspended sediment loads in this area (Eisma & Kalff, 1987; Dyer & Moffat, 1998), and with the idea that sediment loads in the North Sea significantly affect levels of production (Tett et al., 1993; Tett & Walne, 1995). In the German Bight, inputs from the rivers and the atmosphere are high during summer and approach 50% of the available nitrogen pool. However, the high level of production observed here is a combined result of the high level of inputs and high recycling. Light limitation of production by low light levels is thought to be less significant than in English coastal waters. The levels of primary production seen in northwest European shelf seas are maintained by recycling of nutrients during spring and summer. But, what is the net balance between production and respiration? Is the system accumulating organic matter? The international LOICZ programme developed a set of simple mass balance modelling procedures so that the likely degree of accumulation in coastal regions could be assessed in a uniform way world-wide. The LOICZ procedures have been applied in the North Sea by Smith et al. (1997) and in the Irish Sea by Simpson and Rippeth (1998). The procedure assumes that the Redfield ratio can be used to link budgets of C, N and P. The key assumption is that while C and N can be lost from the system as gasses, P is conserved so that the difference between the estimated fluxes of P in and out of the system is a measure of the net balance between production and respiration. Smith et al. (1997) estimated that net ecosystem production is about 2% of primary production. Similar calculations for the Irish Sea found that the uptake of P was statistically insignificant and that the system would appear to be close to trophic balance (Simpson & Rippeth, 1998). It is assumed that riverine inputs of nutrients will be sources to the ocean. That assumption is based on consideration of conditions in other areas of the world ocean with narrow shelves. In such regions, essentially, direct mixing of river and ocean water occurs. In contrast, the area of wide shelf seas makes them significant when the contribution to the oceanic nitrogen budget of denitrification is drawn up. Seitzinger and Giblin (1996) estimated that a net flux from the North Atlantic onto the northwest European shelf of 16 × 1010 mol N y−1 (2440 ktonnes N y−1) is required, in excess of the inputs from rivers and atmosphere, to balance the loss of nitrate caused by denitrification. Similarly on the basis of the LOICZ model, Smith et al. (1997) estimated that in the northern North Sea the loss of N by denitrification exceeded the land and atmospheric inputs to the region. The equivalent rate of denitrification was 0.1 mol N m−2 y−1, and in the southern North Sea it was 0.2 mol N m−2 y−1. In the case of the Irish

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Sea, the estimate was 0.3 mol N m−2 y−1 (Simpson & Rippeth, 1998). These estimates are in line with newer field measurements of denitrification (Lohse et al., 1996; Trimmer et al., 1999). Hydes et al. (1999) used the NERC-NSP data from January 1989 to show that the concentration of P predicted, from a concentration–salinity relationship, at the salinity of source waters at the shelf break, matched the observed concentration. However, the predicted concentration of nitrate was significantly lower than at the shelf break. The nitrate deficit in the southern North Sea was equivalent to a denitrification rate of 0.25 mol N m−2 y−1, assuming a flushing time of one year for the North Sea. These studies suggest that denitrification is a ubiquitous process in shelf seas, and plays a major role in determining the availability of nitrate. In the Irish Sea, data for concentrations of nutrients are available from the CYPRIS station single point time-series since 1955 (Allen et al., 1998). Gowen et al. (2002) compared the observed winter concentrations at CYPRIS with a prediction based on river concentrations and the concentrations in the Antlantic-derived inflow water. Their analysis confirmed a large removal of nitrate. In contrast, observed concentrations of silicate are close to the predicted concentrations. Both Si (Gowen et al., 2002) and P (Simpson & Rippeth, 1998) are conserved within the Irish Sea. However, if the NERC-NSP silicate concentrations from the North Sea in January 1989 are regressed against salinity, the predicted concentration of silicate in the salinity of water at the shelf break is − 0.3 μM Si. This suggests that loss of silicate from the southern North Sea is as great as the loss of nitrate from this system. Silica may be exported as diatom cell wall material. After the spring bloom in the southern North Sea, the tendency will be for this material to be exported across an open boundary in the Jutland current. In contrast, advection from the Irish Sea is restricted by the North Channel (Knight & Howarth, 1999). 9.2.4

Observed distributions of macronutrient concentrations and ratios

The core work of the EU-MAST project NORWESP was the collection of all the available nutrient and productivity data from the region of the northwest European shelf (Radach & Gekeler, 1996). Available time-series data have been analysed by Visser et al. (1996). On the basis of the NERC-NSP data, Prandle et al. (1997) demonstrated that the annual cycle in concentrations can be effectively represented by a cosine wave. The spatially averaged seasonal amplitudes of both nitrate and silicate are approximately equal to their mean values, which is consistent with these being the limiting nutrients. The starting point for production on the shelf is the build up of concentration over winter. The concentration at a particular point is determined by recycling from organic matter in water column and seabed, input from the atmosphere, and exchange with surrounding waters, themselves influenced by oceanic and riverine source waters. In the case of a well-mixed shelf sea in winter, the

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concentration of a nutrient should be predictable from the local salinity, and a knowledge of the concentrations in the riverine inputs and in the appropriate ocean source water, as illustrated by Le Gall et al. (2000). However, the exchange time for mixing of river and ocean water across a continental shelf is proportional to the square of the shelf width divided by the effective diffusivity. This rapid increase with width gives broader shelf seas an interior, across which transports should be considered in distinction to the effectively separate oceanic and coastal boundaries (Huthnance, 1995).

A consequence of the relatively long transport times in northwest European shelf seas is that, in the interior, concentrations cannot be predicted from end member concentrations. As we have seen above, this is particularly true for concentrations of nitrate due to denitrification. In winter, concentrations of nutrients can be lower in shelf seas than in adjacent ocean surface water masses (Hydes et al., 2001; Gowen et al., 2002). Maps of winter distributions of the ratio of N:P show regions of minimum values in those waters furthest from river and ocean source waters (Hydes & Edmunds, 1992; Brockmann & Wegner, 1995; Brockmann & Kattner, 1997). For the southern North Sea, Hydes et al. (1999) report an average N:P ratio at the end of winter (March) of 10 with a corresponding nitrate concentration of 12 μM. During the spring bloom, the N:P ratio rose to 25 and then fell back to below 3 in the summer. This represents an initial limitation of growth by phosphate due to the high ratio of N:P in waters where the dominant source of nutrients is river water. The ratio then fell as production continued using phosphate which was recycled more quickly than nitrate, and nitrate was lost due to denitrification. 9.2.5

Iron

Dissolved iron is operationally defined as that which is detectable in water that has been filtered through a 0.45 micron filter. Increasingly sophisticated methods are to be used to identify the different chemical forms of iron in this so-called dissolved fraction (e.g. Ozturk et al., 2002). However, at the present time there are still many parts of the ocean for which reliable iron data are not available. Johnson et al. (1997) produced the first summary of the variations between oceans of concentrations of dissolved iron. They found that surface water concentrations tended to be <0.2 nmol kg−1 and below a depth of 500 m the average concentration was 0.8 nmol kg−1. The more recent compilation of de Baar and de Jong (2001) suggests that regional variation may be identifiable, but compared to the inter-ocean variation in concentrations of nitrate, the range of variation in concentrations of iron is probably small. Because of a desire not to exclude higher values which may be real, the ranges reported by de Baar and de Jong (2001) are probably wider than that

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which might be considered acceptable in ten years time. They consider that concentrations of dissolved iron in sea water may range over four orders of magnitude – 0.03–1.0nmolkg−1 in surface waters in remote oceans, 0.4–2nmolkg−1 in deep ocean interior waters, 1–10 nmol kg−1 in semi-enclosed and coastal seas, and 1–100nmol kg−1 adjacent to river discharges. Higher surface water values are associated with aeolian inputs and advective transport from continental margins. The concentration of iron associated with particles ranges from equivalent 0.1 to 800 nmol kg−1 from open ocean to coastal seas (de Baar & de Jong, 2001). In most subsurface water, iron is present predominantly in the dissolved phase. The largest source of iron particles is from rivers, most of which is trapped on the shelves. The second largest source is wet and dry deposition from continental aerosols. However, de Baar and de Jong (2001) estimated that flux is an order of magnitude lower than the rate of accumulation of iron in deep-sea sediments. This, taken with the observation that concentrations of dissolved and particulate iron increase towards the margins, led them to suggest that this gradient must be fed from an iron source associated with reductive mobilisation in sediments. Dissolved Fe in oxic seawater is primarily present as ferric iron (Fe-III), and reduced iron (Fe-II) is only present at much lower concentrations (Millero, 1998). The inorganic iron(III) is present as the hydrolysis species (Fe(OH)2+, Fe(OH)3 and Fe(OH)4−). These forms of iron are sparingly soluble, and in the absence of organic chelators and at the pH of sea water, the solubility is 0.1–0.2 nmol kg−1. This is lower than that found in deep ocean waters (Kuma et al., 1996). Therefore, it is considered that the higher concentrations measured in seawater must be stabilised in complexes with organic chelating ligands which themselves are also present at low concentrations in seawater (Gledhill & van den Berg, 1994; Wu & Luther, 1995). The biological uptake of iron is influenced both by iron chemistry in sea water and the cellular uptake mechanisms. The uptake mechanism can vary due to membrane transport of external ferrous or ferric ions in solution, the production of siderophore complexes by the organism which solubilise iron present in bound forms and the ingestion of iron containing particles by myxotrophic plankton. The organic chelation which stabilises iron in solution also means that only about 1% of that present is not complexed. It is thought, therefore, that photoredox cycling may play critical role in making iron available. The cycle consists of the absorption of sunlight by an iron(III) chelate which induces a charge-transfer reaction in which the chelate is oxidised and the iron reduced. The iron(II) dissociates from the degraded ligand, and is then reoxidised and re-complexed. The process is sufficiently slow that it increases the steady state concentration of biologically available iron which can be complexed by iron transport proteins on the cell membrane.

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Plankton biogeochemistry Taxonomy and life forms in the plankton

The marine plankton includes a wide variety of organisms. Microplankton (Dussart, 1965) is made up of essentially single-celled creatures whose population can increase rapidly by asexual multiplication. Zooplankton consists of larger, multicellular creatures which typically make new individuals by sexual reproduction and so have longer inter-generational intervals during which embryos grow into adults. The microplankton includes: (i) the phytoplankton, consisting of cyanobacterial and micro-algal photosynthetic primary producers, (ii) chemo-autotrophic proteobacteria and (iii) micro-heterotrophs, including other proteobacteria, amoeboid, ciliated or flagellated protozoa and fungus-like creatures, feeding on other microplankters or on the organic matter they produce. Copepod crustaceans are the most abundant zooplankters and grazers of microplankton in temperate waters; they are joined by cladoceran and euphausiid crustaceans, and urochordate salps and appendicularians. Higher up the food chain are planktonic carnivores such as chaetognaths (arrow-worms), coelenterate and ctenophore jellyfish, and fish larvae. The taxonomy of micro-organisms (Table 9.2) is in a state of flux, revolutionised in recent years by studies of photosynthetic pigments (Jeffrey & Vesk, 1997) and cellular ultrastucture (Patterson, 1999); by the Serial Endosymbiosis Theory (Margulis, 1993); and by comparison of nucleotide sequences (Wheeler et al., 2000; Patterson & Sogin, 2001; NCBI, 2002). There are now considered to be two prokaryotic domains (Woese et al., 1990), those of the Archaea and the Eubacteria, both without nuclear membranes and cell organelles. The third domain, that of the Eucarya, is characterised by cells with nuclear membranes and other distinct organelles. At the evolutionary base of this domain is the ancestral eukaryotic grade of the Protoctista, formerly seen as one of the five kingdoms of living things (Whittaker, 1969; Margulis, 1993), but now viewed (Patterson & Sogin, 2001) as a collection of lineages, some ancestral to the animal, plant, fungal, brown seaweed and red seaweed kingdoms, and some probably kingdoms in their own right. Unicellular protoctists are also called protists, and include amoeboid, flagellated, ciliated, box-like and hyphal body forms, sometimes assembled into crude colonies. The Protoctista, including photosynthetic algae, phagotrophic protozoa and saprotrophic primitive fungi, is the result of serial symbiosis between an Archaean descendant with a nuclear membrane around its core DNA and one or more types of Eubacteria – in particular, respiring bacteria that became mitochondria and photosynthesising cyanobacteria that became chloroplasts. At the opposite end of the taxonomic hierarchy from domain, kingdom and phylum, is the fundamental category of species, defined (Kelly & McGrath, 1975) as:

Typical marine planktonic forms

Cyanobacteria: PA ‘blue-green algae’, mostly contain phycobiliprotein; filiamentous or otherwise colonial Oscillatoriales† and heterocystous Nostocales† include N-fixers; picoplanktonic members in Chroococcales and Chloroxybacteria (without phycobiliprotein)

Dinoflagellates: 2 dissimilar flagella; many H, probably originally PA by secondary symbiosis with alga. (a) Dinophysidae†, (b) Prorocentrales†, (c) Gymnodiniales†, (d) Gonyaulacales†, (e) Oxyrrhinaceae, (f) Noctilucales, (g) Peridiniales, (h) Pyrocystales

Crytomonads: PA and H flagellates with phycobiliprotein, arisen by symbiosis between H flagellate and red alga

Cryptophyta/ -phycea

Ciliates: bearing cilia (same structure as flagella); oligotrichs (H) with few, elaborate groups of cilia; functionally PA Mesodinium with symbiotic cryptomonads

‘Amoeboid’ protozoa including (phago)H web-feeders, many with symbiotic micro-algae. Actinopoda: Radiolaria with chitinous capsule containing core protoplasm; silicified spines project through pores to support branching pseudopodia. Granuloreticulosea: Foraminifera: foraminifers: with perforated calcarous shell, and unsupported pseudopodia forming a network

Dinophyta/ -phyceae

ALVEOLATA: Ciliophora

SARCODINA

Domain EUCARYA or EUKARYOTA, grade PROTOCTISTA EUGLENOZOA Elastic-bodied uni- or bi-flagellates; Euglenida/ -phyceae: euglenoid PA or H flagellates; Kinetoplastida: Bodonidae: H flagellates

CYANOBACTERIA/ Cyanophyta/ Cyanophyceae

Prokaryote domain (EU) BACTERIA PROTEOBACTERIA Gram-negative bacteria, primitively photosynthetic or chemolithotrophic, with several groups that developed aerobic respiration; original source of mitochondria of eukaryotes; includes anaerobic PA, chemo-autotrophs (e.g. ammonium oxidisers), oxygen or nitrate respiring H pseudomonads

Group

Table 9.2 Microplankton taxonomy and lifeforms

Cryptomonas

Dinophysis, Prorocentrum Gyrodinium, Gonyaulax, Ceratium, Protoperidinium, Noctiluca

Strombidium, Tintinnopsis Mesodinium = Myrionecta

Eutreptiella, Bodo

Trichodesmium, Synechococcus, Prochlorococcus

Nitromonas, Pseudomonas

Example genera

Dictyocha

Green algae: (PA) (a) Volvocales: small flagellates with 1,2,4,8 flagella, some forming flagellated colonies, (b) Chlorococcales: small coccoid cells

Small flagellates (PA) with 1,2,4,8... stiff flagella and organic scales; some colonial forms with roundish, non-flagellated cells in gelatinous matrix and free-living stage

chytrids: hyphal parasites (H) with free-living uniflagellate stage

choanoflagellates, solitary or colonial small (H) unicells with 1 flagellum inside a collar

VIRIDIOPLANTAE: Chlorophyta: Chlorophyceae

Chlorophyta: Prasinophyceae

OPSITHOKONTS: Chytridomycota

Choanoflagellida

Chatonella, Heterosigma

Small, mostly PA, bi-flagellates with specialised haptonema (a) Coccosphaerales, and (b) Isochrysidales, are flagellates, many with a covering of small regular plates (coccoliths) at some stage of life cyle, and/or have spiny organic scales, (c) Pavlovales: flagellates, (d) Prymnesiales†, flagellates, some with aflagellate colonial stage

Haptophyta/ -phyceae or Prymesiophyta/ -phyceae

Monosiga, Diaphanoeca

Pyramimonas, Tetraselmis Halosphaera

Brachiomonas, Dunaliella Nannochloris

Pavlova, Chrysochromulina Emiliana, Phaeocystis

Aureococcus, Aureoumbra

Small or picoplanktonic coccoid algae (PA)

Small bi-flagellates† with many chloroplasts (PA)

Raphidophyta/ -phyceae

Nannochloropsis

Pelagophyceae

Silicoflagellates with many chloroplasts and semi-internal silica skeleton

Very small coccoid algae and small flagellates (PA)

Eustigmatophyta/ -phyceae

Cont

Small PA & H uni- or bi-flagellates including silicoflagellates with silicified scales or cases. (a) Ochromonodales, (b) Paraphysomonadaceae, (c) Synurales

Chrysophyta/ -phyceae

Dictyochophyceae

Cafeteria

Bicoesida: small bi-flagellates; Labyrinthulida: Thraustochytriidae: hyphal parasites with free-living biflagellate

Misc. heterotrophs

Ochromonas, Dinobryon

Asterionella, Chaetoceros Leptocylindrus, Skeletonema Thalassiosira, Pseudo-nitzschia

STRAMENOPHILES (HETEROKONTS): Bacillariophyta/ -phyceae Diatoms: cells with box-like silicified wall, often forming loose chains, almost all PA (by secondary symbiosis with alga), (a) Coscinodiscophyceae, centric diatoms – radially symetrical; (b) Fragilariophyceae – araphid, pennate diatoms; (c) Bacillariophyceae† – raphid, pennate diatoms, bilaterally symmetrical

yes

yes yes

yes

Myxotrophs

hyphal parasites with flagellate stage* amoeboids-pelagic web-spinners*

phagotrophic Dinoflagellates large (Noctiluca) Ciliates (esp. oligotrichs inc. tintinnids) (small, phagotrophic, zoo-) Flagellates*

Heterotrophs

The basis of this table is the taxonomy given by Margulis (1993) which puts most groups at the level of a phylum, with name ending in – [phyt]a) and the algal taxonomy of Tomas (ed.) (1997) (which puts most groups at the level of a class, with name ending in – [phyc]eae). However, the high-level taxonomy has been revised according to the Tree of Life web project (Patterson & Sogin, 2001) and the NCBI taxonomy browser (Wheeler et al., 2000; NCBI, 2002). These are based in cellular ultrastructure and 16s rDNA taxonomy. The Kingdom PROTOCTISTA of the Whittaker 5-Kingdom taxonomy used by Margulis is now seen as a grade containing several lineages or Kingdoms, some of which give rise to the true multicellular organisations of animals, plants and fungi. In the Group column the names of the highest-order taxa, which may be considered Kingdoms analagous to e.g. Kingdom Animalia, are given in CAPITALS; a colon (:) shows hierarchy, a slash (/) separates alternative names. PA = (photo-)autotroph, H = heterotroph, referring to primary features; myxotrophy also common. Groups including toxic members are marked by:†.

(small and picoplanktonic) coccoid cells* filamentous Cyanobacteria (some N-fixing)

(pelagic) Diatoms (centric, raphid, araphid) Tychopelagic diatoms Dinoflagellates . . . large oceanic (Pyrocystis) Mesodinium (small, phyto-) Flagellates* Coccolithophorids Silico Flagellates* flagellates with Colonial stage*

Photo-autotrophs

lifeform terms used in this review excluding Proteobacteria; * indicates lifeform found in several different high-level taxa

Table 9.2 Continued

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a group of similar and related organisms that reacts to, and interacts with, its physical and biological environment in a characteristically unique way [,] . . . a reproductively isolated interbreeding aggregation of populations.

Recognition of micro-algal and protozoan species on the basis of morphology has in recent years been refined by nucleic acid sequencing. But there is a difficulty concerning the part of the definition dealing with characteristic interaction with the environment. The essential feature of such interactions is summarised by Gause’s competitive exclusion principle: to co-exist, species populations must differ in their resource usage (Hastings, 1996). But according to the paradox of the plankton (Hutchinson, 1961; Hobson, 1989; Tett, 1992), there seem to be more species of pelagic micro-organisms than distinct sets of environmental features that allow unique populations of organisms to maintain themselves. At this stage, therefore, any theory of microplankton floristic composition deals (in the main) not with species or even genera, but with lifeforms. An advantage of the lifeform concept is that it is not precisely defined, and so can be applied at any taxonomic level at which it is possible to distinguish kinds of organisms using morphology, life-history and biogeochemistry. In some cases, a lifeform based on biogeochemical or ecological function can include organisms from different taxa. Ryther and Officer (1981) listed seven phytoplankter types which they ranked from the most beneficial centric diatoms, by way of naked and scaled flagellates, green flagellates, pennate diatoms, dinoflagellates and non-motile greens, to the most undesirable bluegreens (i.e. cyanobacteria). Smetacek (1986) noted the category of tychopelagic diatoms for the large heavily silicified centric diatoms of shallow turbulent waters, which are equally capable of living on the sea bed. Riegman (1998) distinguished large diatoms, small diatoms, haptophytes (Prymnesiophyceae, including cocolithophorids and the colonial Phaeocystis), dinoflagellates, mixotrophic algae and cyanobacteria by ecophysiological properties shared with other members of the same taxonomic group. Within each group may evolve specialists in nitrate-, ammonium-, phosphate-, light-limited growth, or with different temperature optima, and this may have led to ecological clusters – i.e. what are here called life-forms – cutting across taxonomic groupings. Reynolds (1987, 1996) used the CSR conceptual model of Grimes (1979) and morphological characteristics to distinguish a variety of life-forms of freshwater phytoplankters, and this approach has recently been applied to marine phytoplankters (Smayda & Reynolds, 2001; Elliott et al., submitted). Smayda and Reynolds describe the three main classes of adaptive strategy: invasive, r-selected, small, fast-growing, high surface-to-volume colonist (C) species; the acquisitive, large, slow-growing but biomass-conserving,

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K-selected, nutrient-stress tolerant (S) species; and the attuning, lightharvesting, attenuated, disturbance-tolerant ruderal (R) species.

Each of their life-forms represents a particular combination of these characteristics. A distinction that clearly cuts across taxonomic groupings within the Protoctista is trophy. Heterotrophs require organic matter both as a source of energy and nutrient elements. Autotrophs use light or chemical energy to synthesise organic matter from inorganic compounds. Myxotrophs (or mixotrophs) can do both; we will use the term to refer, in particular, to organisms that are primarily photo-autotrophs but which use organic matter as a source of nutrient elements. However, it is not suprising that protoctists, given their origin in symbioses between bacterial heterotrophs and autotrophs, exhibit a range of trophies within taxonomic groupings. Any heterotroph which takes particulate food is a phagotroph; saprotrophs use dissolved organic matter, perhaps obtained directly from within other organisms. This discussion of lifeforms amongst heterotrophic as well as autotrophic proticts, leads to the concept of plural microplanktons, the idea that there are persistent associations or reliable correlations between particular autotrophs and heterotrophs. Legendre and Rassoulzadegan (1995) point out that biological oceanographers have generally distinguished between two contrasting trophic pathways in the pelagic environment. [One] goes from large phytoplankton and zooplankton to fish, whereas the [other] comprises small eukaryotic algae and cyanobacteria as well as heterotrophic bacteria and protozoa.

They suggest that there is in reality a continuum of trophic pathways between these extreme systems, and that various points along this continuum are characterised by particular combinations of dominant organisms and nutrient-cycling processes. Lee et al. (2002) have simulated seasonal succession with a model containing two microplanktons; one with diatom-like autotrophs and a small associated population of bacteria and diatom-feeding ciliates and dinoflagellates; and the other with small-flagellate like autotrophs and a larger associated population of bacteria, small zooflagellates and small ciliates feeding on the other organisms. The key biogeochemical distinction is that microplankton 1 requires silica and microplankton 2 does not. 9.3.2

Theories of floristic composition

There are three broad explanations of phytoplankton floristic composition (Tilman et al., 1982). The first utilises differences in the capacity of species or lifeforms to grow in physical environments that differ especially in their vertical mixing intensity. Although nutrients are part of the argument, the explanations focus simply on the supply of the most limiting nutrient. The second type of

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explanation involves the relationship between the ratio of nutrient elements needed for growth and the ambient ratio of these elements. The third deals with variable loss rates due to grazing by protozoans or zooplanktons that preferentially take some species or lifeforms rather than others. No explanation is exclusive and no explanation precludes consequential effects of floristic changes on the consumers of phytoplankton. The three classes of explanation essentially deal with local matters. This is because the abundance of populations of microplankters can potentially change rapidly, at rates of the order of 10−1 d−1, whereas water movement is slow in coastal seas, typically, a few kilometres a day even in regions of relatively high flow (Hill & Simpson, 1988). For example, water flowing through the Sound of Jura in western Scotland seems to adjust its floristic composition quickly to the local intensity of vertical mixing (Jones et al., 1984). Of course, localised explanations beg the questions of where algal seed stocks, nutrient supplies, or long-lived grazer populations originate. Furthermore, if the local factors are properties of moving water, as might be the case for nutrient element ratios, then populations may indeed move with water – as was argued for a 1988 bloom of the toxic flagellate Chrysochromulina polylepis in the Kattegat, Skagerak and Norwegian Coastal Current (Nielsen & Richardson, 1990). Maestrini and Graneli (1991) ascribed it to N enrichment relative to P and Si, coupled with the production of grazer-repellent and algal-supressing toxin by the flagellate. Such conditions may well have been transportable, but Maestrini and Granéli thought that the bloom itself might have originated in several places. 9.3.2.1 Light-nutrient-mixing explanations Light-nutrient-mixing theories for phytoplankton floristic composition are based on the different responses of different life-forms to different physical environments. They stem from the premise that pelagic algae and cyanobacteria need both light and nutrients. However, light penetrates poorly into the sea, so the greatest illumination is to be found near the sea surface. Conversely, gravity sucks particles downwards, so that decay and mineralisation tend to occur predominantly in deeper water. Vertical mixing returns these nutrients to surface waters, but can remove phytoplankters from illumination. Thus, the best places for primary production in temperate shelf seas in summer are often regions where vertical mixing is of intermediate strength, for example at tidal mixing fronts (Pingree et al., 1975; Simpson, 1981). Holligan (1981) reviewed work on plankton distributions across such fronts, concluding that considerable differences between mixed and stratified waters have always been observed . . . Generally, by midsummer, diatoms (including benthic species) tend to be most abundant in more mixed waters [inshore of the front] and dinoflagellates in the front and thermocline [on the offshore side

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of the front], whereas in the surface layer above the thermocline small flagellates, large Ceratium spp. and Rhizosolenia alata are usually the dominant groups.

Lauria et al. (1999) observed how short-term environmental variability allows the co-existence of several algal life-forms in the partially mixed estuary, Southampton Water in southern England. Heavy centric diatoms Coscinodiscus spp. tended to sink from the water column, but were mixed upwards by tidallygenerated turbulence. Periods of stability during tidal slack water allowed surface aggregations of medium sized dinoflagellates Prorocentrum micans and Peridinium trochoideum, which were mixed downwards as turubulence increased during ebb and flood currents. Smetacek (1986) argued that the sinking properties of tychopelagic diatoms such as Coscinodiscus spp. allowed them to remain within estuaries from which they would otherwise be flushed. This argument has been extended to a variety of microplankters that include resting cysts in their life cycle – for example, to the dinoflagellate Lingulodinium machaerophorum, which has cysts that accumulate in the silty sediments towards the head of Loch Creran (Lewis, 1988). Spatial patterns in which distributions of diatoms, microflagellates or dinoflagellates are related to distributions of vertical mixing intensity have also been documented in, for example, Long Island sound (Bowman et al., 1981) and western Scottish waters (Jones & Gowen, 1990). These authors, like Pingree et al. (1978), build on the ideas of Margalef (1978). Margelef considered that phytoplanktonic life forms lay along a continuum from r-selected species, with high growth rates but needing high nutrient concentrations, to K-selected species, which could accumulate biomass slowly but persistently at low nutrient concentrations. Such growth was opposed by losses due to mixing or patch spreading, and so diatoms dominated the spring bloom (high nutrients, high losses) and dinoflagellates dominated summer conditions (low nutrients, low losses). This line of argument is further developed by Smayda and Reynolds (2001) to explain the distribution of nine dinoflagellate life-forms along onshore– offshore nutrient-mixing gradients. Each life-form, which has a particular C-S-R adaptive strategy, plots at different locations in a space mainly defined by axes of (bulk) nutrient accessibility and optical depth, but not excluding the possibility of other dimensions. It is to these other dimensions that we now turn. 9.3.2.2 Biogeochemical controls Biogeochemical explanations of variations in phytoplankton floristic composition work at two levels. The first level is qualitative, and concerns the crude distinctions between algae that require silicon (diatoms and some others) and those that don’t, or between cynaobacteria able to assimilate N2 and all other pelagic photoautotrophs. The second level is quantitative, and concerns the idea that optimum ratios of the nutrient elements required for growth may

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differ amongst species or life-forms. However, the two levels may overlap when discussion concerns ambient nutrient ratios. According to Justic et al. (1995), . . . the atomic Si:N:P ratio of marine diatoms, which are abundant constituents of the coastal phytoplankton, is about 16:16:1, when nutrient levels are sufficient . . . . [However,] . . . rivers historically carried dissolved Si well in excess of dissolved N and P . . . At present many world rivers are beginning to experience a stochiometric nutrient balance (Si:N:P = 16:16:1), or even Si deficiency . . . .

Thus, the decrease in the dissolved Si:N ratio of river water might explain an alleged increase in the frequency of non-diatom blooms in coastal waters. The argument goes back to Officer and Ryther (1980), who point out that it is also necessary to take account of recyling rates, which are faster for P than N and for N than Si. Silicate depleted by uptake during the spring bloom, for example, is often replaced more slowly than DAIN and phosphate, and this may force seasonal succession from diatoms to dinoflagellates or small phytoflagellates. At the second level, phytoplankters may differ in the optimum mix of nutrient elements. Hecky and Kilham (1988) listed optimum N:P atomic ratios ranging from 7:1 to 87:1 for some freshwater and marine phytoplankters, and optimum Si:P ratios ranging from 1:1 to at least 96:1 for freshwater diatoms. Hodgkiss and Ho (1997) argued that most red tide organisms in Hong Kong coastal waters grew best at N:P of between 6:1 and 15:1. Liu et al. (2001) suggested that the low cellular phosphorus requirement of the pelagophyte Aureoumbra lagunensis, as well as the organism’s apparent ability to use forms of phosphorus other than phosphate under severe phosphate deficiency, may partially explain its success in P-limited environments. . . .

Most cyanobacteria, and many cryptomonad algae, use phycobilins as their main photosynthetic accessory pigment. These differ in two ways from the carotenoids used, especially, by the heterokont algae. Phycobilins are haem-like compounds, containing iron, combined with large protein units, hence with a high nitrogen requirement. Carotenoid users may be expected to have a somewhat lower N and Fe requirement than the users of phycobilins. Raven (1998) points out that the picoplanktonic cyanobacterial-descended chloroxybacterium Prochlorococcus, which lacks phycobilins, may make more effective use of iron than the more orthodox picoplanktonic cyanobacterium Synechococcus. Ammonium is held to be preferred by most algae as a source of nitrogen; the uptake of nitrate must be followed by its reduction to ammonium and requires both supply of protons (from photosynthesis or stored energy) and additional iron (Flynn & Hipkin, 1999). In some algae, nitrate uptake is completely

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

inhibited at ammonium concentrations around 1 μM, whereas other species continue to assimilate nitrate when ammonium exceeds 5 μM (Flynn et al., 1999). The former type may be favoured in low-iron environments. Other possible distinguishing characteristics include ability to use inorganic cobalt instead of or as well as organically combined cobalt in the form of vitamin B12 (Granéli & Risinger, 1994) and the ability to use dissolved organic forms of nutrients as well as the inorganic compounds. It might also be that species or life-forms differ in their relative abilities to get different nutrients from their medium: for example, some might invest in machinery for transporting phosphate across the cell wall at the expense of nitrate or ammonium transport systems. Such P. specialists might do better than other algae in low-phosphate environments. Additionally, the ability to assimilate dissolved organic nutrients will confer competitive advantage under conditions of low mineral nutrient concentration – for example, under summer conditions in which messy grazing or defecation liberates N- or P-rich dissolved organic matter (DOM). However, it seems likely that in practice it will be bacterial heterotrophs that acquire most of this DOM, by virtue of their high surface-to-volume ratios. Furthermore, there is evidence that such DOM is in fact rather poor in nutrient content, with a high N:C ratio (Williams, 1995; Sondergaard et al., 2000). The real advantage might thus accrue to phagotrophic myxotrophs that harvest the bacteria, obtaining a ready supply of nutrients at good N:P ratios (see next section) while maintaining a flow of energy from photosynthesis. Many haptophytes have the ability to capture bacteria and small protoctists (Kawachi et al., 1991; Jones et al., 1993), and it may be this that explains the observed ability (Egge, 1998) of Emiliania huxleyi to outcompete diatoms at low ambient concentrations of phosphate. Myxotrophy, however, seems to be a widespread feature of apparently photo-autotrophic protoctists. The primarily photosynthetic armoured dinoflagellate Ceratium furca is able to ingest other protists, including ciliates, which may provide a useful source of N and P when ambient nutrients are scarce (Smalley & Coats, 2002). The photosynthetic flagellate Ochromonas can obtain iron by eating bacteria (Maranger et al., 1998). Novarino et al. report a study of pelagic flagellates (including dinoflagellates) from the south and central North Sea. There were: 44 phototrophs (of which 8 were myxotrophic or possibly so), 35 heterotrophs, and five flagellates whose mode of nutrition is unknown.

9.3.2.3 Ecological controls Verity and Smetacek (1996) have argued that predation or top-down trophic effects are as important as resource-driven or bottom-up factors in structuring planktonic ecosystems. However, grazing control of phytoplankton floristic composition will only occur if secondary consumers are selective for particular algal types. Initial evidence was against such selection. Marshall (1973) reviewed the evidence for dietary restriction in copepods:

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It is clear that copepods can and do select particular foods but they do not do so all the time, and their preferences may change . . . On the whole the gut contents of a filter-feeding copepod are a good reflection of the microplankton present in the sea at the time it was feeding. . . . most copepods are neither purely herbivorous nor purely carnivorous but can change from one mode of feeding to another.

This observation remains valid, as does the view (Parsons & LeBrasseur, 1970; Sheldon et al., 1977) that the main variable in determining what eats what is the relative size of predator and prey. However, there have been major changes in the understanding of trophic relations since the 1970s. Firstly, the diatom-copepod-fish trophic paradigm (Hardy, 1924) has been much modified since the 1970s. It has become clear that small phototrophs often contribute a large share of primary production (Pomeroy, 1974; Joint, 1986); that heterotrophic bacteria salvage much dissolved organic matter (DOM) excreted by phototrophs or spilt during feeding, and return it to the food web by way of the microbial loop (Williams, 1981; Azam etal., 1983); and that heterotrophic protozoa – especially, oligotrich ciliates (Pierce & Turner, 1992), heterotrophic dinoflagellates (Hansen, 1991) and zooflagellates (Fenchel, 1984) – are the main consumers of the smaller microplankters. Although protozoans may compete with larval zooplankters for algal food (Hansen et al., 1994; Gismervik et al., 1996), fully grown copepods are more likely to be predators on ciliates than competitors with them. There is evidence that copepods grow better on a protozoan, rather than an algal diet (Sanders & Wickham, 1993) and that some at least prefer ciliates to diatoms (Verity & Paffenhofer, 1996). A recent study (Irigoien et al., 2000) used removal of taxonomic marker pigments to estimate rates of grazing by Calanus helgolandicus on a variety of phytoplankters during a seasonal cycle in coastal waters of the English Channel. It was concluded that although the copepod seemed restricted to larger algae, it scarcely distinguished between diatoms, dinoflagellates, haptophytes and cryoptophytes. The method could not measure rates of feeding on protozoans, but it was thought that these made an important contribution to diet. Nejstgaard et al. (2001) found that C. helgolandicus preferred ciliates and metazoans to diatoms during a mesocosm experiment in Norwegian waters. We conclude that, in many pelagic environments, protozoans are the main consumers of primary production, with crustaceans feeding mainly on protozoans and not directly on algae. Secondly, theoretical studies of the drag exerted on small organisms and organs as they move through the water, and high-speed photography of copepod feeding, show that true filter-feeding is not possible for copepods (Koehl, 1984). Water is just too sticky a medium on millimetric scales (Lazier & Mann, 1989; Kiørboe, 1993), and so most predators must adopt more raptorial methods to seize or encounter individual prey. Predator morphology in relation to prey

314

BIOGEOCHEMISTRY OF MARINE SYSTEMS

size results in there being an optimum size of prey in relation to predator, but this is no longer the order of magnitude ratio of Sheldon et al. (1977). Instead (Hansen et al., 1994), the linear size ratio between predators and their optimal prey is 1:1 for a dinoflagellate, 3:1 for other flagellates, 8:1 for ciliates, 18:1 for rotifers and copepods, and approximately 50:1 for cladocerans and meroplankton larvae.

where the differences result from differences in feeding organ(elle)s and behaviour. Thus, the ability of dinoflagellates to envelope their prey, or to suck out the contents of a captured cell, allows them to take organisms of size similar to their own. At the other extreme, there is a ratio of order thousands to one between pelagic tunicates which filter water through gills and their prey. Foodweb manipulation experiments demonstrate that doliolid salps and cladoceran crustaceans promote the growth of large algae (by removing smaller phytoplankters), whereas the presence of copepods shifts the size spectrum towards smaller algae (Katechakis et al., 2002). More subtle effects result from the nature of the prey; for, having acquired an item of food, a predator must ingest it. Many life-forms have evolved protection from the siliceous cases and spines of diatoms to the tough-skinned bladder of Phaeocystis that may be difficult to grasp or penetrate, although the thick skin might be more a defence against pathogens than against predators (Hamm et al., 1999). Hansen et al. (1995) hypothesised that Chrysochromulina spp. flourished during a spring bloom in Danish waters because the spines covering their cells made them too large for ciliates to ingest. However, Verity (2000) found in an experimental and modelling study that medium-sized copepods could eat some sizes of colonies of Phaeocystis globosa and Phaeocystis unicells were taken by ciliates. Another form of deterrence might involve the synthesis of toxic compounds, so that a phytoplankter, having been tasted by a predator, may be rejected and survive. It has been argued that production of toxic substances is increased by shortage of one nutrient element in relation to another – for example, by phosphorus deficiency in the haptophyte Chrysochromulina polylepis (Maestrini & Graneli, 1991) and in the saxitoxin-producing dinoflagellate Alexandrium minutum (Maestrini et al., 2000). Maestrine and Granéli argued that the increased toxicity of the haptophyte deterred grazing and was thus one of the factors responsible for the 1988 bloom of this organism in Scandinavian waters. Johansson and Granéli (1999) found that the haemolytic toxicity of C. polylepis was increased by both P deficiency relative to N, and N deficiency relative to P. Bates et al. (1991) found that the neurotoxin domoic acid was produced by the diatom Pseudo-nitzschia multiseries mainly in the stationary phase of the cell cycle, under conditions in which cell division was stopped by lack of dissolved silica, but nitrogen remained available for synthesis of the acid. Maldonado et al. (2002) found that Fe-deficiency and Cu-stress increased production of domoic

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315

acid by P. australis and P. multiseries. They suggested that the acid’s role is not primarily toxic but is to alter the chemical speciation of the metal ions, and thus to increase the ability of the algae to assimilate iron and reduce the harm caused by copper. 9.3.3

Variation in nutrient element ratios and its explanation in terms of plankton biochemistry

Redfield’s (1958) atomic P:N:C ratio of 1:16:106 derives from his and Fleming’s measurements of net-caught plankton. Subsequently, Copin-Montegut and CopinMontegut (1983) surveyed particulates from most of the oceans. Excepting samples with carbonate, they found a similar average C:N:P ratio of 103:16:1. In coastal waters, however, particulates often have C:N higher than the Redfield value of 6.6:1, for example, ranging from 9:1 to 12:1 in the western English Channel in summer (Holligan et al., 1984). Such ratios can be explained by the presence of nutrient-stripped detritus with (atomic) C:N up to 15:1 (Lancelot & Billen, 1985), or above 30 if mineralising bacteria are separated (Verity et al., 2000). The dissolved analogue of detritus also tends to increased C:N ratios, in the range 12–20:1 (Williams, 1995; Sondergaard et al., 2000). In contrast, the C:N ratio of net-separated zooplankton was 5.1:1 in the study of Holligan et al. and Baltic Sea zooplankton had C:N between 5.1:1 and 6.3:1 (Pertola et al., 2002). Table 9.3 summarises estimates of the nutrient element contents, relative to carbon, of a variety of organisms. Gismervik et al. (1996), drawn on for part of this table, point to the existence of two patterns of variation: on the one hand, ‘highly variable stochiometry controlled by growth rates in algae and [heterotrophic] bacteria’ and on the other hand, low variation in elemental ratios in copepods and ciliates. Heterotrophic dinoflagellates and zooflagellates also tend to constant composition (Caron et al., 1990; Davidson et al., 1995). Goldman et al. (1987) placed heterotrophic bacteria in the same category, observing a bacterial C:N:P ratio of 45:9:1 irrespective of food composition. A later study (Goldman & Dennett, 1991) found a similar C:N ratio of 4.5:1 during exponential growth, but lower ratios during the lag phase preceeding this growth. Given their amino acid composition (Parsons et al., 1984), eukaryotic plankters should have proteins with C:N ratios in the range 3–4:1. Compounds with porphyrin rings, such as chlorophyll or the haem cofactors of cytochromes, have C:N between 8 and 14:1. Nucleic acids, the greatest part found in ribsomal RNA and nucleotides, have C:N:P of about 10:4:1. Lipids and carbohydrates are devoid of N and P, except for C:P of at least 40:1 in phospholipid membranes (see Geider and La Roche (2002) for more details). The optimum heterotrophic bacterial C:N:P of 45:9:1 (Goldman et al., 1987; Goldman & Dennett, 1991) indicates cells are largely made up of structural and enzymatic proteins and their cofactors, an N-containing peptidoglycan wall

159–196 196

140–170 110–140–160

copepods and cladocerans net-separated micro- meso-zooplankton ciliates protozoa: zooflagellate (Paraphysomonas), hetero. dinoflagellate (Oxyrrhis) generalised algae algae (various, unspecified) fractionated lake phytopl. large autotrophic oceanic dinoflagellates small flagellates (hapto- and chlorophytes) diatoms freshwater cyanobacteria marine micro-algae (not dinoflag.) pelageophyte (Aureoumbra) marine cryptomonad Cryptomonas marine cyanobacterium Trichodesmium heterotr. marine bacteria ditto

heterotro. bacteria (various)

220

170–200–210

220

2.8 2.7

120 220 184

4.0 2.3

180 190

45 ± 18 81 ± 9 48 43 83

1.4

3.2

175

54 ± 18

3.0 2.8

Qmax/ kQ etc.

129 120–137–155

Qmax

43 43–50–75 14–61 19 ± 7

185–190–200 150–160

151 94–154–200

Redfield copepods (various)

N

Typical Q or optimum

Group/organisms

Nutrient

kQ or Qmin

Table 9.3 Nutrient element ratios in a variety of groups of plankters

Tett and Droop (1988) review Tett and Droop (1988) review Tett and Droop (1988) review (median, n = 17) Liu et al. (2001) converted Sciandra et al. (2000) Sanudo-Wilhelmy et al. (2001) (net-separated chains) Goldman et al. (1987) Goldman and Dennett (1991) (lag phase, exponential growth) Gismervik et al. (1996) review

Tett and Droop (1988) review

Flynn (2001) model Gismervik et al. (1996) review Sommer (1991) Tett and Droop (1988) review (mean and s.d.)

Gismervik et al. (1996) review From Caron et al. (1990) and Davidson et al. (1995)

mg-at (g-at C)−1 Gismervik et al. (1996) review (median and interquartile ranges) Pertola et al. (2002) (Baltic Sea) Holligan et al. (1994) (English Channel)

Reference

Si

P

diatoms in Belgian coastal waters diatoms, marine centric (Skeletonema, Chaetoceros, Thalassiosira)

Redfield copepods (various) copepods and cladocerans protozoa – zooflagellate (Paraphysomonas) generalised algae algae (various, unspecified) large autotr. ocean dinoflag. Pyrocystis small flagellates (euglenoid, hapto- and chlorophytes) diatoms marine micro-algae (not dinoflag.) pelageophyte (Aureoumbra) freshwater cyanobacteria marine cyanobacterium Trichodesmium f.w.cynanobact. Oscillatoria agardhii f.w. pico. cyanobact. Synechococcus heterotr. marine bacteria heterotr. bacteria (in lakes) heterotr. bacteria (in lakes)

23–27–69

6–12–33 1.0

1.7

50–800

22 35–74–124

36

1.3 ± 0.7 1.7 <0.15 9±8

1.3

13

1.7 ± 1.0

10

9.1 7.7–9.7–12

1.8 0.9–1.7–2.5 1.7

9.4 7.0–7.7–11,6 7–24 13.3

2.2–2.5–3.1

6.0

28

7.6

5.1 5.7

Cont

Goldman et al. (1987) Gismervik et al. (1996) from Vadstein (1994) Tett (1998) from Currie and Kalff (1984) mg-at (g-at C)−1 Rousseau et al. (2002) Tett and Droop (1988) review (excluding ‘fast’ Skeletonema) (range and median)

Mastala et al. (1996) converted

Tett and Droop (1988) review Tett and Droop (1988) review (median, n = 9) Liu et al. (2001) converted Tett and Droop (1988) review Sanudo-Wilhelmy et al. (2001) (net-separated chains) Riegman and Mur (1984)

Tett and Droop (1988) review

Flynn (2001) model Gismervik et al. (1996) review Rivkin and Smith (1985), converted

mg-at (g-at C)−1 Gismervik et al. (1996) review Pertola et al. (2002) (Baltic Sea) Tett (1998) from Caron et al. (1990)

0.5

diatom Skeletonema costatum marine cyanobacterium Trichodesmium 37

8

6

1.2

2

48–50

40–110–420

320

Typical Q or optimum

2 1

2–20

55–127

kQ or Qmin

T. weisflogii

generalised alga oceanic diatom Thalassiosira oceanica estuarine diatom: T. pseudonana coastal diatom: T. weisflogii T. weisflogii

plankton

diatoms, freshwater fractionated lake phytoplankton marine diatoms (exponential growth) silicoflagellate (Synura)

Group/organisms

45

107

90

54 30

470

Qmax

90

89

15

27 30

Qmax/ kQ etc.

Sunda et al. (1991) re-analysed Strzepek and Price (2000) Flynn and Hipkin (1999) using Sunda and Huntsman (1997) mean of kQ and Q0 in Flynn and Hipkin (1999) from Harrison and Morel (1986) Droop (1973) converted Sanudo-Wilhelmy et al. (2001) (net-separated chains)

μg-at (g-at C)−1 Bruland et al. (1991) based on Martin and others, 1970s Flynn (2001) model Sunda et al. (1991) re-analysed

Reynolds et al. (2001) model Sommer (1991) Brzezinski (1985) (median and range of values for 27 species) Sandgren et al. (1996) converted

Reference

The table is based on two reviews, supplemented by data (sometimes re-analysed) from research papers. We have attempted to avoid overlaps, except perhaps in the case of the algal data of Gismervik et al. (1996), where the primary sources are not identified. The original data come from (i) experiments with monoclonal populations, and (ii) separated organisms or size fractions from natural plankton. Where the data have been taken from reviews, most are central estimates (means or medians) and ranges (as explained) based on a number of different organisms. Copepod data derived from marine species; other data includes some freshwater as well as marine organisms. Typical Q are observed values for healthy organisms, including both laboratory and natural populations, growing under typical (but often rather undefined) conditions. See text concerning the minimum and maximum nutrient quotas. Quotas given per cell have been converted to per carbon when reliable cell volume data are available, at 10 mmol C cm−3 for diatoms and 15 mmol C cm−3 for others.

Fe

Nutrient

Table 9.3 Continued

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319

and up to 10% of nucleotides. Even at the minimum nutrient content, corresponding to C:N:P of 83:12:1, heterotrophic bacteria have more N and P relative to C than in the Redfield ratio. Although eukaryotes have more nuclear DNA, supported by protein histones, than prokaryotes, the nuclear material is a smaller part of larger cells, and there is presumably room for more carbon-rich energy storage compounds and for phospholipid membranes, explaining the optimum C:N:P of 75:(11–14):1 for the protozoa featured in Table 9.3. Storage of lipids (e.g. Miller et al., 1998) may explain why copepods have typical C:N:P close to the Redfield ratio. Planktonic photoautotrophs display a wider range of cellular composition. For example, the flagellate Pavlova lutheri was grown in one study with C:N:P ratios ranging from 682:66:1 to 88:14:1 (Tett et al., 1985). Similar variability has been shown (Table 9.3) by euglenoids, dinoflagellates, chlorophyceans, cryptomonads, diatoms, pelagophytes, haptophytes and cyanobacteria. The implication is that cells can survive with just a little N and P, grow faster as the cellular nutrient content increases and store nutrient elements taken up in excess of immediate synthetic needs. There are some differences between groups; in particular, phycobiliprotein-containing cyanobacteria have a higher requirement for nitrogen, evidenced by a minimum N content which is higher than that of eukaryotic algae. Such cyanobacteria also have a greater need for Fe (Carr & Wyman, 1986; Brand, 1991). The requirement may be intensified in nitrogen-fixers, which also need Fe as a metallic cofactor for nitrogenase. However, recent calculations (Sanudo-Wilhelmy et al., 2001) suggest that the iron requirement of nitrogenfixing organisms is much more than previously estimated. Iron-containing proteins are essential for photosynthetic and respiratory electron transport. Chlorophyll synthesis is also dependent on iron, and chlorophyll content is diminished under iron limitation. Requirements for iron are also influenced by the nitrogen substrate used to support growth. Both nitrate and nitrite reductases contain iron in their catalytical centres. Iron efficiency models (Raven, 1988) and culture studies (Maldonado & Price, 1999) show that cells using nitrate require 60–70% more cellular iron to support a given growth rate than those assimilating ammonia. An early study of iron limitation in the diatom Skeletonema costatum (Droop, 1973) showed that at least some eukaryote cells can both store a substantial amount of iron and subsist on very little of the metal. Indeed, the lowest content in Droop’s study is close to the theoretical minimum, of order 1 μat Fe:at C, that can be deduced from the photosynthetic requirement of 23 Fe atoms for the cytochromes in a photosynthetic unit (PSU) of 900 molecules of chlorophyll (Flynn & Hipkin, 1999). It has been claimed that oceanic algae require less iron in relation to phosphorus (Brand, 1991) or carbon (Sunda et al., 1991) than do coastal algae. Diatoms vary in their wall thickness and hence their need for silica in relation to other elements. Brzezinski (1985) gives a range of values for marine diatom

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

Si:C from 0.04 to 0.42 with a median of 0.11. Freshwater diatoms tend to be thicker walled, as exemplified by the standard Si:C ratio of 0.32 used in the model of Ryenolds et al. (2001). Silicoflagellates’ silica requirements can be as high, as evidenced by the colonial freshwater flagellate Synura petersenii, which has silicified scales (Sandgren et al., 1996). However, the flagellate can become temporarily scale-free. Similar quantitative studies of marine silicoflagellates have not been reported, although there have been observations of skeleton-free Dictyocha speculum (Henriksen et al., 1993).

9.3.4

Quantitative theory for nutrient element ratios

The relationship between nutrients and the growth of populations of microorganism can be described in three ways. The simplest theory is the one developed for organic-carbon-limited bacterial growth by Monod (1942) and popularised for application to phytoplankton by Dugdale (1967). In this MONOD theory, the rate of uptake of dissolved nutrient (per unit biomass) depends on ambient concentration S: S u = u max ⋅ ------------ks + S and growth rate,

μ=u ⋅ q

(9.1)

–1

The crucial parameters are the concentration kS at which uptake is half of the maximum rate and the fixed ratio (or yield) q−1 at which nutrient is converted to biomass. The yield may be the Redfield ratio or some other optimum composition. Tilman et al. (1982) used the model to show how freshwater phytoplankters of different optimum composition or different half-saturation concentrations, might succeed to different extents depending on the ambient ratios of nutrient elements. Although the assumption of constant yield may be appropriate for pelagic heterotrophs, it is now seen to be too simple for accurate prediction of the growth of phytoplankters (Droop, 1983; Sommer, 1991; Ducobu et al., 1998). The second level, or cell-quota theory, allows organisms to vary their content of nutrient Q (and hence their yield of biomass from assimilated nutrient). It is increasingly referred to as the DROOP model after one of its authors (Droop, 1968, 1983). In principle, the quota should be defined as the ratio of nutrient to biomass (Droop, 1979) and will here be understood as the population (atomic) ratio of the nutrient element to carbon. A simplified version of the theory and some deductions from it, is given in Box 1 (refer page 348). The key equation (ignoring physical transports) is:

INFLUENCE OF NUTRIENT BIOGEOCHEMISTRY ON ECOLOGY

dQ ------- = u – μ ⋅ Q dt

321

(9.2)

Q Where growth rate, μ = μ max ⋅ ⎛ 1 – ρ . ⎛ -----⎞ ⎞ ⎝ ⎝ k Q⎠ ⎠ The quota can vary between a minimum content kQ and a maximum Qmax, and ρ is 1 for a limiting nutrient and greater than 1 otherwise. The ratio Qmax/kQ determines a population’s storage capacity for a given nutrient. The ratio kQj1/kQj2 helps to determine relative limitation by nutrients j1 and j2. For example, Sj1/Sj2 < kQi, j1/kQi, j2 may indicate that nutrient j1 limits species i, although as Box 1 shows, other factors also play a part. Riegman (1998) has argued that we should expect the biochemistry of major taxa to determine their basic characteristics, whilst individual species can evolve into particular ecological niches by changing some of their nutrient uptake or growth parameters. Cell-quota parameters, in particular kQ (Table 9.3), have been estimated for enough algae and cyanobacteria to allow this hypothesis to be tested by comparisons between groups or lifeforms, and by seeking differences amongst species within some taxonomic groups. However, there are complications. Adaptation can be physiological and reversible, as in the case of laboratory algal populations that have been shown to switch between fast- and slow-forms (Harrison et al., 1976). Techniques for estimating subsistence quotas include regression of steady-state growth rates on population nutrient content and the measurement of final ratios in batch cultures. In addition to the practical difficulties in growing algae or cyanobacteria under controlled, ideally axenic conditions, there can be statistical problems in comparing results gained by different techniques. Furthermore, a population may be thought to be limited by one nutrient when it is in fact limited by another (or light), the outcome being that the subsistence quota estimated for the claimed limiting nutrient is in fact ρ · kQ and so higher than the true value. A final problem is that nutrient quotas are often reported per cell rather than per unit biomass or carbon. It is of course possible to convert to a carbon basis when information is given about cell size, but this introduces additional variability and perhaps substantial errors where size is known under conditions other than those of zero growth. Published values for subistence and maximum quotas that are reported in units of atoms nutrient (at C)−1 or can be reliably converted are given in Table 9.3. Table 9.4, which summarises ratios of subsistence quotas, draws on additional papers. Given such interpretational difficulties, which add errors to those given in the tables, only a few generalisations can be reliably drawn: (1) silicon subsistence quotas are variable amongst diatom species, with typical freshwater and tychopelagic marine diatoms having heavier walls

M M M

F F F

M M M

M M

F M M

M M M

Micro-algae Small flagellates Centric diatoms

Cylindrospermopsis raciborskii Cyanobacteria Cyanobacteria

Alexandrium catanella Gymnodinium mikimotoi Heterocapsa circularisquama

Pavlova lutheri Aureoumbra lagunensis

Asterionella formosa Skeletonema costatum Diatoms

oceanic phytoplankters coastal phytoplankters Synechococcus, 5 strains

Cyanophyceae

Bacillariophyta, Fragilariophyceae Coscinodiscophycea Bacillariophyta

Dinophyceae Dinophyceae Dinophyceae mean (of Dinophyceae) Haptophyceae Pelagophyceae

Cyanophyceae Cyanophyceae Cyanophyceae

as below Chlorophyceae, Haptophyceae Bacillariophyta

kQ ratio (2)

m-at Fe:at P <0.1 0.8–10 2–40

at Si:at N 3.3 0.13–0.52

40 133

22 13 12

36

at N:at P

0.56

16

13 54

40 27

Brand (1991) note (5) Brand (1995) note (5) Brand (1995) note (5)

Tett and Droop (1988) note (3) Tett and Droop (1988) note (4) Tett and Droop (1988) re-analysed for group medians

Terry et al. (1985) Liu et al. (2002)

Matsuda et al. (1999) Yamaguchi and Itakura (1999) Yamaguchi et al. (2001)

Istvanovics et al. (2000) Tett and Droop (1988) reanalysed for medians using Tett and Droop (1988) median NkQ and other values in Table 3.2 for PkQ

from Tett and Droop (1988) medians in Table 3.2. Tett and Droop (1988) re-analysed for group medians Tett and Droop (1988) re-analysed for group medians

(1) F = freshwater, M = marine. (2) unless stated, the values are the ratios of cellular subsistence quotas (usually elemental mass or g-atoms per cell) estimated in the study given in the reference; values ascribed to the review by Tett and Droop (1988) are based on subsistence quotas calculated per unit atom of organic carbon and may have come from the earlier review by Shuter (1978); the values for nitrogen and phosphorus do not necessarily refer to the same set of organisms or studies. (3) Nitrogen subsistence quota taken as 43 atoms N (atom C)−1 based on Shuter (1978); the silicon quota was based on Tilman and Kilham (1976), using another source for cell volume. (4) Nitrogen susistence quota taken as 44 atoms N (atom C)−1, based on Shuter (1978); the silicon quotas were for the ‘fast’ (low silicon) and ‘slow’ (high silicon) strains studied by Harrison et al. (1976). (5) estimated from the ambient Fe:phosphate ratio at which nutrient limitation of final cell yield shifts from one nutrient to another in batch cultures (bactericised?) of (22 overall) species of marine phytoplankters.

(1)

Organism

Table 9.4 Values of ratios of subsistence quotas

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(2) (3)

(4) (5) (6)

323

than pelagic marine centric diatoms such as Skeletonema, Thalassiosira and Chaetoceros with a characteristic subsistence value of 0.03 at Si:at C; the ratio of Qmax/kQ for N and Si is lower (2–4) than for P and Fe (5–90); phyocobiliprotein-containing cyanobacteria and cryptomonads have higher nitrogen subsistence quotas (0.08 at N:at C) than other photoautotrophs; they may have higher FekQ; typical (eukaryote) photo-autotrophs with carotenoid accessory pigments have N kQ of 0.05 at N:at C; large oceanic dinoflagellates have lower nitrogen subsistence quotas (0.02 at N:at C) than other phytoplankters; although there may be specific exceptions, typical phosphorus subsistence quotas for all groups of phytoplankters appear to be within the range 0.001–0.002 at P:at C.

Combining points (4) and (5) implies that typical eukaryotic autotrophs require a minimum of 1–2% of their carbon in nucleic acids and about 20% in proteins; see also Geider and La Roche (2002). The implication of point (2) is that phytoplankters can often draw on a larger cellular reserve of phosphorus than of nitrogen (Fig. 9.1). This is part of the explanation for the tendency of marine phytoplankton to be N limited rather than P limited. Additionally, since a typical value of the ratio NkQ/PkQ is 28 at N (at P)−1, a phytoplankter N:P ratio with the Redfield value of 16:1 implies N limitation, the value of NQRedfield/ NkQ being less than that of PQRedfield/ PkQ. However, a phytoplankter with a non-standard kQN/ kQP may become P limited at a different internal or ambient N:P ratio. For example, Table 9.4 suggests that the mean NkQ/ PkQ for coastal red tide dinoflagellates might at 16:1 be somewhat lower than that for most algae, whereas the small pelageophyte Aureoumbra lagunensis has a high ratio of 133:1. This high ratio is the result of a very low PkQ, implying a nucleic acid content relative to biomass that is less than a tenth of that in most eukaryotes. Brand (1991) estimated the subsistence optimum Fe:P molar ratios of cultivated dinoflagellates, flagellates and picoplanktonic cyanobacteria of the genus Synechococcus, from the initial ambient Fe:PO4 ratio at which nutrient limitation of final cell yield shifts from one nutrient to the other in batch experiments. In the absence of significant carry-over of nutrient by the populations used to start the experiments, it should be true that if FeBmax/PBmax = 1 then Fe St = 0/PSt = 0 = FekQ/PkQ (see Box 1, Equation 9.6). The tested strains came from coastal waters and the ocean. Irrespective of their origin, Synechococcus needed more iron relative to phosphorus than did eukaryotic algae. It also seemed that the subsistence Fe:P ratio was less in oceanic eukaryotes, especially coccolithophorids, than in coastal eukaryotes. This distinction in relation to origin is supported by the finding (Sunda et al., 1991) that the coastal diatom Thalassiosira weisflogii had a greater FekQ than its oceanic relative T. oceanica. However, Droop (1973) had earlier reported a subsistence quota for the coastal

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

1.00

µ /µ max

0.50

0.00 0

10

Q /kQ

20

30

1.00 f inhib(Q )

0.50

QP QN

0.00 0

10

Q /kQ

20

30

200 mmol N (mol C)–1

50 N

Redfield P

1.5 mmol P (mol C)–1

33 0

kQ /kQ

30

16 10

molar N:P ratio

7

20

30

Q /kQ

Fig. 9.1 Cell-quota theory for control of photo-autotroph growth by internal nitrogen or phosphorus. Q is the cell quota for the nutrient, in atoms of the element per atom of organic carbon. kQ is the minimum value, or subsistence quota. μ/μmax gives growth as a proportion of maximum rate. The function finhib(Q) multiplies nutrient uptake (which is also a function of ambient concentration) and brings it towards zero as Q tends towards Qmax. The third part of the diagram compares typical ranges of values of cellular N and P content and show how these contribute to variation in the cell N:P ratio.

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325

diatom Skeletonema costatum, that was lower than the values for Thalassiosira, and this single counter-example casts some doubt on generalisations. Finally, cell-quota theory treats the entire cellular content of a nutrient as being the pool controlling growth rate (under limiting conditions), and deals with multiple nutrient interactions empirically. At the third level of description, MECHANISTIC models aim to embody realistic accounts of the main biochemical processes and pools within cells. A recent example (Flynn & Hipkin, 1999; Flynn, 2001) deals with nitrogen, phosphorus, silicon and iron as well as the carbon content of cells, photosynthesis, and the uptake competition between ammonium and nitrate. However, the model embodies many parameters and there is currently insufficient information to use it to distinguish between groups or species of phytoplankters. Its characteristic nutrient quota parameters are included in Table 9.3 except for those for silicon, which are cell-based. 9.3.5

Differences in abilities to assimilate different nutrients

For most nutrients and algae, the ratio umax/μ · Q is so much greater than 1 when S/kS > 1 that, differences in the ability to assimilate nutrient can have little bearing on the relative success of species or life forms. However, at low nutrient concentrations, S/kS << 1 and uptake properties are important. The most obvious outcome under such conditions is the dominance of small organisms because, all other things being equal, they have greater umax as a result of larger surface-to-volume ratio and less intense diffusional restriction on nutrients. According to Sunda and Huntsman (1995), [e]xperiments with coastal and oceanic phytoplankton clones representing different algal groups and cell sizes indicate that cellular iron uptake rates are similar among the species when rates are normalized to cell surface area.

Morel et al. (1991) give an equation for growth rate in a steady state at low external concentration of a trace metal such as iron. Slightly modified, it is:

Where

S′ ⋅ K L ⋅ L T μ = ------------------------Q u max k L ⋅ L T ≅ ---------ks

(9.3)

S′ is the ambient concentration of available metal, kL is the rate constant for the uptake reaction in which the metal combines with a transport ligand, whose abundance per cell is given by LT. Having drawn attention to the likelihood that the Fe-transport molecules of eukaryotic algae are membrane bound siderophores, like those of cyanobacteria, Morel et al. conclude that there is:

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

. . . no direct evidence demonstrating that the specific growth rates of phytoplankton in oceanic waters are limited by trace metals or for that matter by any major nutrients. In the high-nutrient, low-Fe waters . . . experimental data simply show that Fe additions increase the biomass in incubation bottles after a few days. . . . in these waters . . . the indigenous community is growing quickly and . . . this growth is supported chiefly by NH4+ as the N source . . . According to the results of diffusion calculations, to double once a day at an ambient NH4+ concentration of 0.3 μM, cells >~7-μm radius must utilize NO3− to supplement the available NH3. Because growth on NO3− presumably requires additional cellular Fe . . . [for use in nitrate reductase], large cells . . . should be Fe limited by virtue of both their size and their need for NO3−.

Experiments in which the ocean itself, or ocean water in bottles, has been enriched with iron, confirm this conclusion; the enrichment stimulates an increased biomass of large algae, especially lightly silicified, raphid pennate diatoms (Martin et al., 1991; Hutchins et al., 2002). Such a finding is not entirely academic in respect of the coastal seas of northwest Europe. It is presumably because of the relative availability of iron here that centric diatoms are able to flourish and can be stimulated by macronutrient enrichment up to limits set by the availability of silicate. Should iron become less available – for example, because of complexing by a land-derived binding agent (Hutchins et al., 1999) – dominance might be expected to shift to ammonium-users. Timmermans etal. (1994) showed that iron depleted cultures of prymnesiophytes or prasinophytes have less nitrate reductase than iron-enriched cultures, and that iron enrichment during bioassays of low-iron northern North Sea water, enhanced nitrate reductase activity. The ability to convert di-nitrogen into organic nitrogen is a property of certain large cyanobacteria, either members of the Nostocales, with heterocysts, or those such as Trichodesmium in which chains clump together to create the low-oxygen environment in which nitrogenase can function. It is not altogether clear why marine pelagic nitrogen-fixers should be found only in the low-salinity waters of the Baltic Sea and in some tropical oceans. It may be that only these environments have persistent conditions of low turbulent shear coupled with high irradiance, which enables local oxygen depletion to take place in bright light. Atlantic ocean Trichodesmium seem to be limited by ambient P rather than Fe (SanudoWilhelmy et al., 2001), despite the iron requirement of the N-fixing enzymes. Whatever constraints exist on the biogeochemical composition of phytoplankters, uptake abilities ought to be potentially more variable. It should be possible for the transport sites for one nutrient in the cell wall to be increased at the expense of sites for other nutrients. However, this does not seem to occur in the case of iron (Sunda & Huntsman, 1995), perhaps because of evolutionary pressures that have pushed iron uptake in all species toward the maximum limits imposed by diffusion and ligand exchange kinetics.

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327

There may be inter-specific differences in relative ability to assimilate ammonium and nitrate, as exemplified for the raphidophyte Heterosigma carterae and the coccolithophorid Emiliana huxleyi (Flynn et al., 1999). Apart from this, the bulk of variation between these phytoplankters in their ability to take up dissolved nutrients seems explicable by size differences and interactions relating to size, as exemplified above for iron, and we will not further pursue this theme. However, uptake of dissolved inorganic compounds across the cell wall is not the only way to get scarce nutrients. Many eukaryotic algae (but no diatoms) are myxotrophs, with the ability to ingest particles (see Section 9.3.2) or even colloidal iron (Nodwell & Price, 2001). Non-living particles may carry iron or phosphorus, but the ideal source is small microplankters that have already harvested the nutrients from low external concentration. Turning from myoxtrophs to pure heterotrophs, we see that although bacteria require a higher content of nutrient elements than they usually get from their food, they can use the energy source provided by organic carbon to assimilate nitrate, ammonium, phosphate or labile dissolved organic nutrients (Middelburg & Nieuwenhuize, 2000; Rodrigues & Williams, 2001); failing these, they mineralise relatively more carbon or the excess nutrient element (Goldman & Dennett, 1991). Protozoans are somewhat less demanding of their food C:N:P, but are not able to supplement this from inorganic nutrients. Zooplankton can often optimise their nutrient element input by feeding preferentially on protozoans instead of phytoplankton. 9.3.6

Theoretical conclusions

The preceding subsections suggested that the responses of photo-autotrophs to ambient nutrient ratios are largely determined by (a) their relative subsistence quotas, and (b) their uptake abilities, influenced by size, growth rate and phagotrophic capacity. Heterotrophs are buffered from ambient ratios by the intervention of autotrophs, and tend to an optimum composition. Nutrient element needs and assimilation abilities of different groups of plankters are summarised in Table 9.5 which draws on, but goes beyond the evidence in Tables 9.3 and 9.4. Its contents may be taken as hypotheses for predicting how plankton and phytoplankton in particular, will respond to particular ambient ratios. 9.4 9.4.1

Effects of ambient nutrient ratios on plankton Introduction

The previous section considered the composition of plankters and their ability to assimilate nutrients in theory and in single-species populations growing under controlled conditions. The present section considers the extent to which the hypotheses of Section 9.3 are supported by data for real plankton, made up of different species growing under complex and varying conditions.

50:–:1 30?:–:1 50:–:1 40:–:1

40?:–:1 16?:–:1 40?:(11):1 16:–:1 27:14:1

A

A A Am, H Am, H

A Am Am Am, H, Hm A A H

Cyanobacteria – chain-forming

Cyanobacteria – picoplanktonic, phycobiliprotein Cyanobacteria – picoplanktonic chloroxybacteria Cryptomonads – small flagellates Most other non-silicified small flagellates – Euglenozoa, Heterokonts (Bicoesida, Chrysophyceae, Haptophyceae, Viridioplantae), Opisthokonts Gelatinous colonial stages of some of above flagellates, e.g. haptophyte – Phaeocystis Raphidophytes – small flagellates Silicoflagellates – Heterokonts (Chrysophceae, Dictyophyceae) Alveolata – Ciliophora (ciliates), Dinophyta (dinoflagellates) Pelagic diatoms – Bacillariophyta

Tychopelagic diatoms – Bacillariophyta

Pelagic crustaceans (copepods, euphausiids etc.)

(1) N:Si:P

Proteobacteria

++ 16:1

phagotrophy, vertical migration may exploit nutrient gradients, toxicity poor nutrient uptake because of size; slow sinking poor nutrient uptake because of size; fast sinking phagotrophy

+ ++

phagotrophy, toxicity ability to lose silica scales or skeleton

+ + 12:1

difficult to predate

? 12:1

Notes wide range of biogeochemical abilities; H are saprotrophs N-fixation, some toxicity, buoyancy-driven vertical movement may exploit P gradients small size, good nutrient uptake small size, good nutrient uptake phagotrophy phagotrophy, some parsitism (in fungal types), some toxicity

+

++ + ++ +

+++

(3) N:P 10:1

(2) Fe +?

A = autotrophs, mostly photo-autotrophs; Am = photo-autotrophs with some (myxotrophic) ability to obtain and use dissolved or particulate organics; H = heterotrophs; Hm = heterotrophs with autotroph symbionts, thus also myxotrophs. (1) This column gives the typical atomic ratios of subsistence quotas in the autotrophs or myxo-autotrophs; raphidophyte ratios are based on those cited by Yamaguchi et al. (2001); (2) This column indicates iron need in autotrophs and myxo-autotrophs, on the basis that phycobiliproteins require additional iron, as does N-fixation or nitrate-reduction; (3) This column gives optimal atomic ratios for the heterotrophs.

27:75:1

(50):–:1

?

Contains (A), H

Group or lifeform

Table 9.5 Conclusions about nutrient element ratios in marine plankters

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329

The evidence that we present here comes from several types of source. The most clear-cut are the results of well-controlled experiments in which mesocosms are initiated with natural microplanktons and then enriched at various ratios of the nutrient elements. The most convincing evidence comes from long-term studies of phytoplankton in waters that have become nutrient enriched during the study, with associated changes in nutrient element ratios. There are too few of these, and we focus on one, the Helgoland time-series. Finally, supporting evidence comes from observations of phytoplankton seasonal successions at selected sites (with different ambient nutrient element ratios) in northwest European seas. Oceanographers customarily distinguish between new and recycled nutrients (Dugdale & Goering, 1967; Eppley & Peterson, 1979). However, the equation of new with nitrate and recycled with ammonium needs modification. For present purposes, recycled nutrients are those that become available within the photic zone due to local grazing and mineralisation. New nutrients are (i) those found in the sea in late winter, before the vernal growth of phytoplankton and (ii) those supplied to the photic zone from outside during the productive season. In the case of a mesocosm, new nutrients are (i) those initially present and (ii) those supplied during an experiment. In the sea, the resupply during the growth season may come from diapycnal mixing, upwelling, estuarine circulation, or human-enriched inputs from air, rivers, or direct discharges. It is generally possible to make a good estimate of the ambient ratios for new nutrients, either from winter or initial nutrient concentrations, or from supply concentrations. It is much harder to do so for recycled nutrients. Pools are low and turn over rapidly, and any case may be by-passed by myxotrophy. 9.4.2

Time series: Helgoland and the German Bight

Europe has few sites at which hydrography, nutrients and phytoplankton have been measured simultaneously and intensively, during a sufficiently long period, to demonstrate progressive changes in conditions against a background of interannual variability. One of these sites is close to the island of Helgoland in the inner German Bight of the North Sea. As a result of anthropogenic enrichment of water from the Elbe and other rivers discharging into the Bight (Radach, 1992; Hickel et al., 1993), winter N:Si ratios here, which were in the range 1–2 in the late 1960s, increased to 4–8 in the mid-1990s (Radach et al., 1990). The ratio change is the result of increasing DAIN and decreasing silica. Changes in N:P have been more complex, with winter values in the range 16–32 in the early 1960s falling into the range 4–16 in the 1970s and increasing again to 16–32 in the mid-1980s. According to Hickel et al. (1993), starting in: the early sixties, phosphate concentrations rose for about a decade, levelling off to about twice the former concentrations for another decade, and then decreasing (since 1982) as a result of phosphate-reducing measures. Nitrate

330

BIOGEOCHEMISTRY OF MARINE SYSTEMS

concentrations, however, have only increased since 1980/81, following Elbe river flood events. In 1987, three times the former concentrations were reached.

The nutrient changes have been associated both with an increase in the total stock of phytoplankton (Gillbricht, 1988; Radach et al., 1990; Hickel et al., 1993) (Fig. 3.2, Gillbricht, 1988) and with changes in the seasonal succession. In the 1960s, diatom abundance peaked in July, followed rapidly by a peak in flagellates reaching almost the same biomass. In the early 1980s, diatoms peaked in April at nearly three times the old biomass, but the flagellate peak remained in August, also reaching nearly 3 times its previous level. The flagellates included dinoflagellates, especially species of Ceratium and small naked forms (Radach et al., 1990). According to Table 9.5, a typical pelagic diatom NkQ/ SikQ ≈2. Their Si-limited growth should consume N:Si in the ratio ρ NkQ/SikQ – i.e. at least twice as much DAIN as silicate. Thus a winter ambient N:Si much above 2 should result in a diatom bloom followed by blooms of other algae, which seems borne out by the Helgoland observations in the early 1980s. However, explaining the changes between the 1960s and the 1980s as a sole result of nutrient enrichment is open to several objections. The phytoplankton at Helgoland may have been participating (Hickel, 1998) in long-term changes affecting the North Sea as a whole, and resulting from climate change as much as nutrient enrichment (Tett & Mills, 1991; Reid etal., 2001a,b). There is much short-term variability at this location amidst strong gradients and flows, subject to year-to-year variability in weather and river discharge. Hickel et al. (1993) points out that, in the control of phytoplankton growth, hydrographical factors possibly dominate. Additional nutrient input by Elbe river floods did not always result in elevated phytoplankton stocks near Helgoland, while extended periods of vertical density stratification of the German Bight water caused large plankton blooms.

Nevertheless, we think that such variability in effect can be discounted by systematic analysis of sufficient data, as demonstrated by the methods used for time-series analysis by Radach et al. (1990) and the use of salinity as a normalising variable by Gillbricht (1988). Salinity normalised analysis showed that moderate levels of ambient nutrients remained during summer in the 1980s. The persistence of silicate might explain why diatoms continue to be abundant, if not dominant, during summer but neither why the peak of diatoms, a summer phenomenon in the 1960s, has shifted to spring (Fig. 9.2), nor why overwintering flagellate biomass was so high in the early 1980s, despite the plausible supposition that winter phytoplankton is limited by light and not nutrients. It is possible that some of flagellates were phagotrophs (Hickel et al., 1993), exploiting a food chain originating in stored or imported organic matter and thus able to flourish during winter.

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331

μg C dm–3 120 diatoms 80–85 diatoms 62–67 80

40

120 flagellates 80–85 flagellates 62–67

80

40

0

0

3

6 month

9

12

120

μg flagellate C dm–3

62–67 80–85

80

40 April 0

0 April

40

80

μg diatom C dm–3

120

Fig. 9.2 Phytoplankton changes at Helgoland (redrawn from Gillbricht, 1988). Flagellates include dinoflagellates and small flagellates, and some may be heterotrophic. The values are monthly means over periods of 6 years of samples taken at least twice-weekly.

Attempts to model the phytoplankton changes include those of Dippner (1998) who was able to simulate observed diatom and flagellate changes using a Monod growth model with a fixed ratio of C:Si:P of 106:16:1, forced with a combination of phosphate increase and silicate decrease. The ecosystem model ERSEM II (with nutrient kinetics that uncoupled uptake of N, P and Si) was used (Pätsch & Radach, 1997) to hindcast the development of the ecosystem of the North Sea during the years 1955 to 1993. . . . The comparison of the hindcast with the long-term

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

observations at two sites in the continental coastal zone of the North Sea shows that the long-term behaviour of phosphate, nitrate and silicate is simulated well. . . . The flagellates at Helgoland, however, experience much more pronounced annual cycles with much less interannual variability in the hindcast than in the observations.

9.4.3

Mesocosm and other competition experiments

A difficulty with conclusions based on observations is that correlation of changes does not prove causation. As Gillbricht (1988) pointed out for the Helgoland data set, the correlation between phytoplankton and [nutrient] need not have anything to do with causality; exceptional algal blooms have been observed and reported in the literature since in the 19th century.

More rigor is provided by mesocosm experiments, in which large volumes of seawater are held in floating plastic bags and subjected to various treatments. An early example is the final CEPEX experiment commenced in July 1978 in Saanich Inlet, Canada (Grice et al., 1980). Two bags were enriched with phosphate and nitrate, and the second of these was further enriched with silicate and partly vertically mixed. Diatoms increased initally in both bags, but continued to flourish only in the second. Experiments with deep bags in the eutrophic Seton Akai of Japan showed that dissolved silica depletion leads to shifts of dominant species from larger to smaller diatoms or flagellates (Harada et al., 1996). Table 9.6 summarises the results of some recent experiments in northwest European waters, using somewhat smaller bags than those in the Japanese and the Canadian experiments and hence allowing a greater variety of treatments, with replication. It is not easy to see an overall pattern in these results, given the variety of initial conditions as well as treatments. Our table emphasises on nutrient ratios at the start of what in some cases were batch cultures and semicontinuously diluted cultures in other. In experiments carried out by Gilpin et al. (2003) and Roberts et al. (submitted) in the Trondheimsfjord, high initial N:Si ratios had two effects (Fig. 9.3). The first was depletion of silica before nitrate. The second was relatively greater abundance of dinoflagellates and flagellates relative to diatoms once the silicate had run out, although in all cases diatoms dominated biomass throughout the 16-day experiments. Taking account of studies reported by Escaravage et al. (1999), Williams and Egge (1998), Egge and Jacobsen (1997) and Jacobsen et al. (1995), it seems that at N:Si ratios of 2 or less, diatoms remain important, whereas ratios of 4:1 and more always give way to non-siliceous algae – although the change may take some time.

May 94

Land-based with water from Oosterschedlt, Netherlands

July 97

May 94

April 92

Raunefjorden, Norway (Bergen MBS)

Raunefjorden, Norway (Bergen MBS)

Raunefjorden, Norway background Si throughout

June 94

June 99 and 2000

Start month(s)

Trondheimsfjord (Trondheims LSF) phosphate in excess

Site of mesocosm and initial water and plankton

Table 9.6 Mesocosm experiments

2 + sea 2

n = 10:1 16:5 16:1

nd nd nd

2

2 2 2 2 2 2 3 1 1 2 2

5

6

n (b)

7:1 10:1 17:1 8:1 11:1 18:1 n = 4:1 15:1 15:1 15:1 15:1

<6:1

4:1 1.0:1 1.3:1 2.2:1 1.7:1 2.3:1 3.7:1 n = 1:1 1.5:1 >15:1 3:1 15:1

<3:1

N:P (a)

1:1

Treatment N:Si (a)

control (treatments compared with this) diatoms up to 20x; E. huxleyi up to 3x; Phaeocystis cells/colonies up to >20x diatoms up to 10x; E. huxleyi up to 10x; Phaeocystis up to >20x;

Diatoms dominate, nitrate depleted before silicate Diatoms dominate, but relatively more flagellates; silicate depleted before nitrate diatoms dominant (changing species), Phaeocystis cells and colonies important ditto, Phaeocystis relatively most abundant diatoms dominant (changing species), Phaeocystis cells and colonies important ditto, Phaeocystis relatively most abundant fjord and controls: diatoms and flagellates diatoms maintain dominance Emiliana huxleyi becomes dominant diatoms dominant, some Phaeocystis more E. huxleyi, Phaeocystis disappeared

Main consequences

Cont

Egge and Heimdal (1994)

Williams and Egge (1998); Sondergaard et al. (2000); Pond et al. (1998) Egge and Jacobsen (1997); Nejstgaard et al. (1997)

Escarvage et al. (1999)

Gilpin et al. (2003); Roberts et al. (submitted)

Reference

April 88 June, July 88

Raunefjorden, Norway (Bergen MBS)

n = 10:1 16:1

n = 4:1 16:1

4:1 4:1 4:1

16:0.2

nd

1.5:1 1:1 >6:1*

N:P (a)

Treatment N:Si (a)

1 1+1 1

2

n (b) diatoms <10x; E. huxleyi up to 10x; Phaeocystis little increase control: diatoms to unidentified flagellates, cryptophytes and prasinophytes enriched: diatoms to Phaeocystis – flagellate, also high choanoflagellates and ‘microzooplankton’ diatoms to Phaeocystis diatoms dominate throughout coccolithophorid dominance, then Phaeocystis

Main consequences

Egge and Aksnes (1992)

Jacobsen et al. (1995)

Reference

(a) Atomic ratio of ambient nutrients, which in most cases were enriched. Based on initial ratios in the case of short experiments, otherwise on resupply ratio in case of enrichment or observed concentration when no enrichment. n = maintained at natural levels. * marks nutrient that was not enriched, i.e. remaining at local sea concentrations or with recycling within bag. (b) number of replicates of treatment.

Feb. 91

Start month(s)

Raunefjorden, Norway (Bergen MBS) background Si throughout

Site of mesocosm and initial water and plankton

Table 9.6 Continued

INFLUENCE OF NUTRIENT BIOGEOCHEMISTRY ON ECOLOGY

(a)

335

30.0

μM nitrate 25.0 N:Si = 1:1 N:Si = 4:1

20.0 15.0 10.0 5.0

μ M dissolved silica 0.0 0.0

4.0

102

N:Si = 1:1 N:Si = 4:1

8.0

12.0

(b)

log10(x 10–4 μ m3 L–1) flagellate biovolume 101

log10(x 10–4 μ m3 L–1) diatom biovolume 100

100

101

102

103

Fig. 9.3 Scatter plots of results from 16 day experiments in mesocosms (Gilpin et al. (2003) and Roberts et al. (submitted)) treated with nitrate and silicate in 1:1 and 4:1 molar ratios. (a) 1999 and 2000 nutrient data. Note exhaustion of nitrate before silicate at 1:1 and of silicate before nitrate at 4:1. (b) 2000 phytoplankton biovolume data: flagellates = autotrophic dinoflagellates plus autotrophic nanoflagellates. Note greater relative abundance of flagellates compared to diatoms at 4:1 towards the end of the experiments, when the highest biovolumes occurred. We are grateful to K. Davidson (SAMS-UHI), L. Gilpin (Napier) and E. Roberts (Swansea) for permission to use these data.

Some of the authors of the studies interpreted their results in relation to ambient levels of a particular nutrient rather than to a particular initial or supply ratio, in effect emphasising kS(μ) rather than kQ. Thus Egge and Aksnes (1992) found that diatom dominance occurred irrespective of season if silicate concentration exceeded a threshold of approximately 2 μM. Flagellate dominance changed

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BIOGEOCHEMISTRY OF MARINE SYSTEMS

to diatom dominance within a few days after nutrient addition resulting in silicate concentrations above this threshold. Dominance of Phaeocystis sp. appeared on several occasions after the bloom of another species, but never at high silicate concentrations. The success of the diatom group seemed to be due to a high inherent growth rate at non-limiting silicate concentrations.

Additionally, it seemed to Egge (1998) that diatoms were not able to dominate when phosphate was deficient, although silicate and nitrate were in excess. From a dominance of 72 and 74%, for the entire experimental period, in enclosures supplied with all three nutrients, the diatom dominance was reduced to 45 and 55% in enclosures where phosphate was not supplied. Growth of the flagellate community appears to be less affected by low phosphate concentration.

The flagellates that grew instead of diatoms included unicellular and colonial Phaeocystis or the coccolithophorid Emiliana huxleyi under N and P enrichment (Egge & Aksnes, 1992; Egge & Heimdal, 1994; Jacobsen et al., 1995; Egge & Jacobsen, 1997; Williams & Egge, 1998) and unidentified flagellates, cryptophytes and prasinophytes (Jacobsen et al., 1995) under non-enriched conditions. Egge and Heimdal (1994) concluded that E. huxleyi grew well at low concentrations of phosphate, and was a better competitor than Phaeocystis sp. as long as the sea water temperature was about 10°C or higher and the surface irradiance was 20mol m−2 day−1 or higher. Phaeocystis sp. had a higher phosphate demand, but was a better competitor than E. huxleyi at sea water temperatures lower than 10°C and at surface irradiance lower than 20 mol m−2 day−1, if nutrients were available.

Aksnes et al. (1994) made a model of diatoms, E. huxleyi and other flagellates, potentially limited by N, P and Si. The cocolithophorid was given a higher growth affinity for phosphate than the other two algae. Simulations were compared with results of mesocosm experiments. They could not account for the high observed numbers of E. huxleyi in terms of the . . . orthophosphate in these experiments. . . . By assuming an organic phosphorus source that adds phosphorus to the orthophosphate pool, however, the fit between the model and the observations [was] improved.

But it also seems possible that the coccolithophorids were getting phosphorus by myxotrophy. Mesocosm experiments start with an inoculum of plankton that is only partly under experimental control. An alternative approach is to mix single-species populations. Riegman et al. (1996) combined up to 18 laboratory-cultivated

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species under discontinuously diluted conditions with various ratios of nutrients (nitrate, ammonium and phosphate, plus excess silica when diatoms were involved). They concluded In comparison with the other non-diatoms, the mixotrophic [prymnesiophyte] C. [hrysochromulina] polylepis grew fast under all tested nutrient and light limitations. E. huxleyi grew well under nitrogen (N) limitation (with nitrate as N source) and at irradiance levels from 15 up to 500 μmol quanta m−2 s−1. No growth of calcifying cells could be detected under N limitation when ammonium was used as N source. [The cryptomonad] Rhodomonas grew reasonably well under ammonium-N limitation and grew fast at the highest irradiance. The dinoflagellates were poor competitors compared to the Prymnesiophyceae.

Riegman et al. (1996) argued that the prymnesiophytes excelled in these artificial conditions by virtue of their high asexual cell division rate, whereas the dinoflagellates would do better in natural habitats for which they are particularly fitted by specific life-cycle and behavioural characteristics (such as cyst formation or vertical migration). Sommer (1994) performed competition experiments with up to eleven species of marine phytoplankton at various ratios of silicate:nitrate and various light intensities. Diatoms became dominant at Si:N ratios > 2.5:1 while flagellates were the superior competitors at lower ratios. The light supply did not influence the competitive position of diatoms and nonsiliceous flagellates in general, while ii: was important in determining the outcome of competition at the species level. In the 11 species experiments, Stephanopyxis palmeriana was the dominant diatom at high light intensities. It shared dominance with Lauderia annulata at medium and low light intensities and high Si:N ratios. Pseudo-nitzschia pungens was the dominant diatom at low light intensities and relatively low Si:N ratios. The green alga Dunaliella tertiolecta was the dominant flagellate at high light intensities, while at low light intensities the prymnesiophycean Chrysochromulina polylepis and the cryptophyte Rhodomonas sp. were also important.

9.4.4

Observations at sea

Table 9.7 collects data on nutrient ratios and dominant phytoplankters for a number of sites in northwestern European waters. It is not comprehensive, being biased towards sites with which the authors are familiar in reality or through the literature. However, it does comprehend a range of nutrient ratios that is sufficient to test hypotheses about impacts on phytoplankton. The sites in the Swedish Himmer fjord and in Dutch coastal waters provide time-series analagous to those from Helgoland. The Himmer fjord exchanges with the Baltic, so has low salinity. Much of its anthropogenic nutrient comes from a waste water treatment plant, which has in most years been used to strip

N

e

E

E

E

N

e

T

M

MiH

MiH

Hdc

Hdc

Hdc

western central North Sea, 1980s East Anglian coastal waters, late 1980s Netherlands coastal waters, c. 1990 Helgoland early 1980s, at 32 psu (inner) Himmerfjord, Sweden, Baltic: 6 psu Loch Creran, Scotland, 1970s Loch Striven, Scotland, 1980s–90s

N

T

18, 12, 1.5 (16, 17)

6, 6, 0.5 (14)

28, 25, 1.0 (12)

46, 11, 1.5 (10)

30, 10, 0.9 (7)

4.6, 3.0, 0.7 (6) 4–6, 4–5, 0.6 (1) 14, 5, 0.9 (6)

8–10, 4–5, 0.7 (1)

1.5:1

1.0:1

1.1:1

4.1:1

3:1

2.7:1

<1.5:1

2.0:1

12:1

12:1

28:1

31:1

33:1

16:1

6–10:1

13:1

Winter Winter Hydro (a) Nutr (b) Winter μM [N] [Si] [P] (c) N:Si (d) N:P (d)

northern North Sea, 1970s–90s

Site

Table 9.7 Ambient nutrient element ratios and phytoplankton succession

diatoms (17)

diatoms, Phaeocystis (7) diatoms (small flag.) (10, 11) diatoms and dinoflag. (13) diatoms (15)

Phaeocystis (6, 7)

diatoms (6)

diatoms (2)

Spring bloom (e)

18:5:1 (10) 75:19:1 (12) 2.5:7:1 (16) 7:7:1 (17)

20:8:1 (8)

R

R

Resupply N:Si:P (f)

small flagellates, dinoflagellates, diatoms (7, 9) flagellates (diatoms and dinoflag. esp. Ceratium) (11) dinoflag. and cryptomonads, diatoms and Mesodinium (13) diatoms (small dinoflag. and small flagellates) (15, 16) diatoms and/or small dinoflag. and small flag., diatoms or dinoflag. (17)

small flag. inc. coccolithophorids (dinoflag. esp. Ceratium) (3, 4, 5) small flag. (dinoflag. esp. Ceratium) (3, 6) tychopelagic diatoms (6, 7)

Summer (e)

E

Irish Sea – Liverpool Bay (England)

29, 4.9, 1.7 (18)

9.5, 5, 0.8 (18)

6:1

1.9:1

17:1

12:1

small flag. esp. Phaeocystis, diatoms (18) Phaeocystis colonies, diatoms (18) 17:2.5:1 (18)

67:33:1 (18)

diatoms (18)

diatoms, small dinoflag. (18)

(a) Hydrographic status: M = mixed, MiH = mixed, intermittent haline stratification, T = seasonal thermal stratification (winter mixed), H = year-round, predominantly haline stratification (with dc = weak density-forced circulation or DC = strong ditto). (b) Nutrient supply status: N = natural, e = somewhat anthropogenically enriched, E = strongly anthropogenically enriched. (c) Winter inorganic nutrient concentrations for observed winter maximum, whenever that occurs. (d) (Atomic) ratios calculated from preceding (concentration) column. (e) Group(s) dominating the autotroph biomass. A,B means type A followed by type B; A(B) means A dominant but B important; (f) resupply ratio only given if there is a significant input of new nutrient, either natural or anthropogenic; calculation requires assumptions about sources. R = reccyling system. References: (1) Radach and Gekeler (1996); (2) Gillbricht in Radach (1980); (3) Reid et al. (1990); (4) Head et al. (1998), Widdicombe et al. (2002); (5) Lee et al. (2002); (6) Mills et al. (1994), Gieskes and Kraay (1984); (7) Tett et al. (1993) and other NSP data; (8) Colijn, ICES report, 1991; (9) Justic et al. (1995) for Rhine; (10) Gillbricht (1988) – resupply ratios are observed summer ratios at 32 psu; (11) Radach et al. (1990); (12) Tett et al.; (13) Elmgren and Larsson (1997) also assumption that summer Si comes solely from outside (14) Jones (1979); (15) Tett et al. (1981, 1985) and unpublished; (16) Tett and Edwards (2003); (17) Tett et al. (1986, 2001) (18) Gowen et al. (2000) – Phaeocystis classified as ‘microflagellate’ but may have been colonial, especially in Liverpool Bay.

M

e

Irish Sea – Irish coastal MiH

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P more stringently than N in order to ensure P limitation in the fjord (Plaza et al., 1990; Engqvist, 1996; Elmgren & Larsson, 1997) and hence avoid the N-fixing cyanobacterial blooms which now typify the open waters of the Baltic during fine summers (Finni et al., 2001). The Netherlands waters include the Wadden Sea, in which Phaeocystis colonies and large diatoms increased in several steps between 1974 and 1994 (Philippart et al., 2000), and offshore waters (de Vries et al., 1998), in which nutrient and chlorophyll concentrations show gradients of up to one order of magnitude perpendicular to the coast within the first 30–50 km offshore. Time-series analysis reveals significant decreasing trends for dissolved inorganic phosphorus (40%) and total phosphorus (35%) and an increase in the dissolved inorganic N:P ratio from 25–30 to 40–55 over the period 1988–1995. Trends in nitrogen (−15%), silicate (stable), and chlorophyll are smaller and generally not statistically significant. The trends in phosphorus reflect a proportional and immediate response to decreasing riverine inputs.

The northern North Sea is strongly influenced by inflow from the North Atlantic, which has a naturally high N:Si ratio; the ratio falls towards the central North Sea, presumably as a result of the denitrification discussed in Section 9.2. Conditions in the western Irish Sea are also largely set by Atlantic inflow, in this case to the south of Ireland. Nutrients in Liverpool Bay, part of the eastern Irish Sea, are enriched by the discharge of the river Mersey, which also creates intermittent haline stratification. The two Scottish fjords are the sea-loch Creran, in near-pristine condition with high Si in the 1970s, and loch Striven, enriched with anthropogenic nutrient from western Scotland by way of the rivers discharging into the inner Firth of Clyde. Two substantial points emerge from the table. One is negative: variations in N:P ratio seem to have no discernable effect. The positive point concerns diatoms, which stand out as the characteristic life-form and group of northwest European waters, almost irrespective of N:Si ratios. The obvious explanation would seem to be that diatoms are equipped, by virtue of the ability to grow fast at low illuminations, to dominate the spring phytoplankton increase. In waters where there is re-supply of silicon, either by local mineralisation or external input, diatoms remain important, often intermittently throughout summer, at least to the extent of forming an autumn bloom. In strongly stratified waters, such as those of the central and northern North Sea, local recycling resupplies some N and P to the photic zone allowing growth of probably myoxtrophic small flagellates and vertically migrating Ceratium spp. When initial N is high, relative to Si, utilisation of initial silica may leave an excess of N for use by Phaeocystis or other flagellates. This is generally considered to be the explanation for the nuisance blooms of Phaeocystis colonies along the continental coast of the southern North Sea (Lancelot et al., 1987; Lancelot, 1990). However, other factors may be involved. For example, during

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1988–89, Phaeocystis made up a larger share of the spring phytoplankton in East Anglian waters than in those off the Belgian and Dutch coasts or in the German Bight, although N:Si ratios were higher in these continental waters (Tett etal., 1993). The importance of Phaeocystis before the main spring diatom bloom in Irish coastal waters, where the N:Si is at 1.9:1, relatively low, supports the idea that the alga is most adapted to physical or optical conditions in shallow, stirred waters with a high suspended load rather than being a particular beneficiary of high N:Si. Finally, Croot et al. (2002) made a summer transect of the Skagerrak. In waters of Baltic origin, with elevated trace metals levels, but very low macronutrients, a mix of dinoflagellates and haptophytes dominated the low biomass. In the Jutland current, which had high dissolved iron concentrations, a mixed bloom . . . of diatoms (major species – Leptocylindricus danica) and dinoflagellates (Ceratium sp.) was present. In the waters of the central Skagerrak derived from the North Sea, below the low salinity Baltic water, a large diatom (major species – L. danica) bloom . . . was present . . . below the pycnocline and was located at the nutricline for silicate. The lowest concentrations of trace metals were found in the water of North Sea origin. Synechococcus-like cyanobacteria were observed in the upper waters across the survey area.

Coincidence does not confirm causality, of course, but these findings are, at least, compatable with the hypotheses discussed previously. 9.5 9.5.1

Discussion and conclusions Introduction

Do, or can, shifts in ambient nutrient ratio, perturb the balance of organisms in the plankton? That was the question asked at the start of this review. We have considered: (Section 9.2) the controls on ambient ratios; (Section 9.3) theory based on knowledge of the biochemical composition of plankton and the results of laboratory experiments with single strains of pelagic phototrophs; and (Section 9.4) observed relationships, in the sea and mesocosms, between microplankton floristic composition and nutrient ratios. The theory suggests that particular ambient nutrient ratios should favour particular phytoplankters. Much speculative explanation has been mined from this rich seam, but in practice it is difficult to identify nutrient-ratio effects on primary producers with certainty against a background of hydrographical and ecological variability. Only when such variability is eliminated or controlled, as during mesocosm experiments, does a clear biogeochemical signal begin to emerge. Even in these cases it is not always clear-cut. Our objective in this section is to reach a set of reliable conclusions about nutrient ratio effects and to use these as a basis for proposing limits for ambient ratios in northwest European seas that should prevent undesirable changes to the balance of organisms from this cause alone.

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Organisms themselves do not experience ambient nutrient ratios directly. The ratios are modified by assimilatory and internal storage mechanisms. Thus, there are especial difficulties in relating floristic composition to ambient ratios under lownutrient, rapid-cycling, summer conditions. For the sake of simplicity, we therefore emphasise the ratio of elements in new nutrients. In the case of the temperate seas of northwest Europe, these ratios are winter values and those in external enrichments including river discharges and experimental additions to mesocosms. The next 3 sub-sections consider nutrient ratio effects pairwise, starting with N:Si and proceeding by way of N:P to P:Fe. We end by considering the extent of flexibility in the Redfield ratios and concluding in a Panglossian fashion that the micropelagic ecosystems of northwest European seas are actually rather resistant to nutrient-ratio induced disturbance to the balance of organisms. 9.5.2

Do high ambient N:Si ratios favour flagellates?

Officer and Ryther (1980) suggested that increased fluxes of N and P relative to Si would, given especially the slower recycling of Si, tend to shift the balance of organisms from diatoms to flagellates. This was because diatoms use silicon in approximately a one-to-one atomic ratio wtih nitrogen (Redfield et al., 1963). The flagellates associated with coastal eutrophication need only nitrogen and phosphorus. . . .

Justic et al. (1995) also took 1:1 as the standard N:Si ratio for diatoms and saw higher ratios leading to non-diatoms blooms. The evidence from northwest European seas and mesocosms certainly supports the general argument of both sets of authors. However, the view that the critical N:Si ratio is 1:1 needs revision. Both theory and observation suggest that some diatoms can make a little silicate go a long way, and that the N:Si ratio must at least exceed 2:1 and perhaps 4:1, for the floristic balance to be perturbed away from diatoms. As Officer and Ryther pointed out, high ambient N:Si does not supress diatoms, but initially stimulates diatom growth, leading to exhaustion of dissolved silica while DAIN and DAIP remain. Relatively high ambient N:Si ratios (2:1 or more) occur naturally in the Atlantic inflow to the northern North Sea and the western Scottish shelf. Grazing and sedimentation of spring bloom diatoms remove silicon effectively from the euphotic zone of the northern North Sea, and the well-stratified waters of summer are dominated by small phytoflagellates, including coccolithophorids and large phototrophic (possibly myxotrophic) dinoflagellates. Further from the Atlantic inflow, N:Si ratios diminish, probably because of denitrification, aided on the west coast of Scotland by the run-off of N-poor, silica-rich water from the acidic highlands. In the central North Sea, the succession from diatoms to flagellates is little changed, but in the fjordic sea-lochs of western Scotland diatom dominance may endure throughout the summer as a result of continued resupply of silica by runoff as well as that brought from deeper water by estuarine circulation.

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High N:Si ratios also now prevail in some of the coastal waters of northwest Europe. These waters include those between northeastern France and southern Norway, impacted by the discharge of large rivers such as the Rhine and the Elbe and the outflow from the Baltic Sea. They also include the parts of the Celtic and Irish Seas affected by discharges from the rivers Severn and Mersey. In all cases, the rivers themselves carry abnormally high nitrogen loads as a result not only of human enrichment but also of the canalisation and cleaning of estuaries (Billen et al., 1985; Billen et al., 1991) and relatively low silicon loads for the reasons discussed by Conley et al. (1993). Near Helgoland, the changes have increased the summer dominance of flagellates. In the stratified Kattegat and Skagerak, the result is (probably) an increase in flagellate blooms during late spring and summer. In the vertically mixed, enriched, waters on the coast of the Netherlands, Belgium and East Anglia, and in the similar waters of western Britain, the high N:Si ratio is widely held to have resulted in increased blooms of colonial Phaeocystis, the argument being that silicate exhaustion by diatoms leaves DAIN (and phosphate) to be exploited by non-siliceous algae. However, the hypothesis of a strict causal relationship is somewhat undermined by evidence discussed in Section 9.4, including the observation that Phaeocystis is important during the spring in shallow and relatively weakly enriched waters to the east of Ireland as well as in the well-enriched waters of Liverpool Bay. Certain physical conditions may be as much a pre-condition for Phaeocystis as particular nutrient ratios. These conditions may be the alternation of mixed and weakly stratified conditions that occur as a result of tidal straining in shallow and often well-stirred waters receiving a large freshwater inflow. The other characteristic of these waters is a high load of suspended sediments. Why are diatoms the characteristic pelagic algae of northwest European seas when they are subject to several handicaps? Heavy cell walls cause them to sink, and so they fall out of the photic zone under conditions of low turbulence. Exhaustion of dissolved silica prevents their further growth. They compensate with good growth efficiency under low light conditions (Tett, 1990) but why don’t they adopt the strategy of some silicoflagellates of dispensing with their silicified parts when necessary? Conversely, why are silicoflagellates less common than diatoms in the sea, given this trick and, in some cases, the ability to get nutrients by phagotrophy? It may be that diatoms have a collective trick: as Si levels fall, there may be a succession from species with a high Si requirement to species with a low requirement, e.g. Rouseau et al. (2002). The record seems to be held by a fast strain of Skeletonema costatum, with a subsistence quota of 0.006 at Si:at C (Harrison et al., 1976). This is about a quarter of the typical marine pelagic diatom subsistence quota of Table 9.3, and only about a twentieth of diatom median silica content according to Brzezinski (1985). Finally, there is some evidence that the main nutrient constraint on diatoms themelves is not the N:Si ratio but the ambient concentration of dissolved silica. As noted in laboratory-cultivated Thalassoisira (Paasche, 1973) and in mesocosms

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(Egge & Aksnes, 1992), growth ceases when this concentration falls below a threshold, in the range 0.3–2μM. The explanation may be that, in many diatoms, cell-wall construction draws directly on ambient dissolved silica. If this is inadequate, cell-wall construction ceases and hence the cell division cycle is halted. This argument is not incompatable with that involving ambient ratios of new N:Si, whereby diatoms grow until they use up all Si, leaving N if the initial ratio was sufficiently high and/or when recycling occurs. The only difference is that diatoms cease growth before all Si is exhausted. As Fig. 4.2, which combines results from a number of mesocosm experiments, demonstrates, silica use is sometimes almost complete, and at other times incomplete as required by a threshold model of use. Clearly, there is more to find out about diatoms and silica. 9.5.3

Do non-Redfield ambient N:P ratios perturb pelagic ecosystems?

The idea that phytoplankton take up nutrients in the Redfield ratio has led to the deduction that deviations of the ambient N:P ratio outside the range of 10–16:1 should disturb the floristic balance of the phytoplankton. However, assimilatory ratios of N:P can diverge greatly from Redfield, because of luxury uptake, and many algae can store P against future need. An argument from subsistence quotas suggests that ultimate growth limiting ratios of N:P are typically closer to 30:1 than 16:1 for most algae. The exception might be some dinoflagellates and raphidophytes, where the kQ ratios are close to 16:1. This lends some support to the hypothesis of Hodgkiss and Ho (1997), who argued that decreasing N:P (from 20:1 to 11:1) in Hong Kong waters made them a better environment for red tide causative organisms, which had a relatively low optimal N:P. Some mesocosm results have suggested that high N:P (or phosphate deficiency) favours the coccolithophorid Emiliana huxleyi, which grew best, compared to diatoms and Phaeocystis, at N:P of 80:1 (Egge & Heimdal, 1994). However, E. huxlyei’s myxotrophy might help it under these conditions. Mesocosm experiments have also been run with very low ambient ratios of N:P. Down to 3:1, there seems little obvious perturbation other than that explicable by changes in N:Si. Low N:P ought to favour cyanobacteria where these can fluorish, as in the Baltic Sea, because they can fix nitrogen to match excess phosphorus. Loads of N and P into the Baltic Sea are estimated to have increased, respectively, fourfold and eightfold since the 1950s (Larrson et al., 1985), with a corresponding increase of blooms of toxic, filiamentous, N-fixing cyanobacteria (Finni et al., 2001). The external loading ratio for N:P in the Baltic proper is now at least 40:1, which should lead to P limitation. However, it has been suggested that much of the sea is, in fact, primarily N limited and that the success of cyanobacteria in summer is a result of the spring bloom’s exhaustion of DAIN, whilst leaving measurable phosphate (Elmgren & Larrson, 2001). That is to say, the argument concerns succession (and, implicitly, the existing of summer-stratified conditions) as much as initial or loading elemental ratios. However, the argument

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needs a little more examination, because algal capability for luxury uptake of phosphate might be expected to be removed in ratio ρ PkQ/NkQ > 1/27–40 (using ratio data from Table 9.5) and hence to deplete phosphate completely during the spring bloom, given an initial N:P ratio of 40:1. Hence the issue may turn on the relative rates of mineralisation of N and P. 9.5.4

The possibility of iron limitation in shelf seas

The link between iron and the growth of oceanic phytoplankters was put forward in the 1930s (Gran, 1931), but it was not until the 1980s that developments in clean sampling and analytical methodology allowed the first reliable, accurate and precise measurements of iron in the sea, e.g. Martin and Gordon (1988). The 1990s saw a rapid expansion of our knowledge of the chemistry of iron in sea water, its availability to plankters, and the consequences for both plankton and biogeochemical cycling (e.g. Raven, 1990; Greene etal., 1991; Millero, 1998; Turner & Hunter, 2001). Much of this interest has been driven by a desire to understand the apparent lack of productivity in the HNLC (high nitrate low chlorophyll) regions of the ocean and the link with atmospheric CO2 and climate suggested by Martin (1990). Marginal and shelf sea have tended to be discounted as likely regions in which trace metals are biolimiting. Martin (1990) thought that meeting biological iron requirements is relatively easy in neritic waters, [where] resuspended bottom sediments and associated iron rich oxides, colloids etc., occur together with elevated concentrations of iron. Hence excess nitrate is never observed in coastal upwelling environments such as those off the west coasts of Africa and North and South America.

However, although evidence is growing that plankton can assimilate iron by a variety of pathways (Sunda, 2001), because these specific abilities are specific to different organisms and because organisms differ in their iron requirements, there is increasing suspicion that limitation by trace metal deficiency may be a widespread phenomenon. It has been demonstrated that growth limitation by iron does occur in coastal upwelling regions (Bruland et al., 2001). Much of the work reported to date has focussed on off-shelf environments (de Baar & de Jong, 2001). However Gledhill et al. (1998) reported decreasing levels of reactive iron as a bloom progressed, indicating that iron was not present in great excess in the northern North Sea. Also in the northern North Sea, Timmermans et al. (1998) found that although ‘no signs of true iron deficiency were found. . . . Strong support for the effect of Fe on cell physiology is given by the increase in the f-ratio’. A recent study in a Norwegian Fjord found that the speciation of iron may be critical (Ozturk et al., 2002). 9.5.5

Trophic consequences of ratio changes – a Panglossian conclusion?

The concept of biome is widely used in terrestrial plant biogeography. A biome is a distinctive combination of plants and animals, characterised by

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a uniform life-form of vegetation. Centric diatoms would seem to be the characteristic life-form of the pelagic biome in northwest European seas, in which the metaphorical understorey is of small phytoflagellates, which can fluorish when the main vegetative cover is cleared. Unlike the temperate deciduous woodland which is the natural biome of much of northwest Europe’s land, the centric diatoms of the spring bloom can, however, largely vanish after making their major contribution to the year’s primary production, to be succeeded in waters made temporarily oligotrophic by summer stratification, by low biomasses of phytoflagellates and phototrophic dinoflagellates. Our conclusion in this article is that the centric diatom dominance of the biome is more resistant to major disturbance than is sometimes thought, except in regions where physical conditions already favour other life-forms, in the way that conditions in the oligotrophic ocean gyres select for picophytoplankters. In our temperate shelf seas, there would appear to be two such types of region. The first is the shallow, tidally stirred, turbid, intermittentely weakly stratified waters that seem to favour Phaeocystis colonies, a predisposition enhanced by enrichment at high N:Si. The second is the summer-stratified, low-salinity, waters of the Baltic Sea, where N-fixing cyanobacteria are favoured, especially during fine summers, by enrichment at low N:P. In all other cases, serious deviation of ambient nutrient ratios from the flexible Redfield values discussed below, may stimulate the flagellate and dinoflagellate understorey, and the impact of this disturbance on the rest of the marine ecosystem is likely to be enhanced because of greater nutrient-induced abundance of the flagellates. However, the actual impact may be less than anticipated. Diatoms persist, as seen at Helgoland, despite big increases in N:Si. Flagellates, dinoflagellates and colonial Pheocystis appear to take over dominance only where physical conditions already predispose. Despite earlier plausible arguments concerning ecosystem flips as a result of nutrient enrichment and changes in ratios (Tett & Mills, 1991), recent evidence concerning the palatability of different algal groups is increasingly equivocal (Gismervik et al., 1996). Indeed, the evidence begins to suggest that zooplankton may get better nourishment from microplankton dominated by myxotrophic algae and heterotrophic protozoans than from a diatom-dominated bloom. The argument that diatom blooms provide a better supply of sinking food than do flagellates is contradicted by some of the mesocosm work. On the one hand, Wassmann et al. (1996) found that both primary production and sedimentation were higher in a mesocosm enriched with N, P and Si, in which diatoms flourished. On the other hand, Svensen et al. (2001) reported that addition of dissolved silica to mesocosms caused higher primary production and a shift from a flagellate to diatomdominated phytoplankton community. However, contrary to expectations, sedimentation of chl a was lower . . . where diatoms dominated than where flagellates prevailed . . .

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The pelagic food web seems pretty good at coping with variation in nutrient element ratios. Primary producers tolerate wide discrepancies from the optimum by adjusting their growth rate to match the availability of limiting nutrient, and storing any excess of others. Myxotrophs supplement their uptake abilities by capturing smaller organisms. Microplankton heterotrophs convert their food input into protoplasm that is rich in N and P by burning off the excess carbon or other nutrient. Some bacteria are especially good at this, taking detritus with C:N ratios that may be as high as 30:1 and converting it into bacteria with C:N of 5:1 – albeit at very low efficiency (Verity et al., 2000). Finally, zooplankters preferentially predate protozoans and so gain food already close to their Redfield ratio optimum composition. 9.5.6

A flexible Redfield ratio?

As we have discussed, the sum of trophic processes does often lead to average and mesozooplankton C:N:P ratios close to those of Redfield, but these processes can start from ambient ratios that are quite different from Redfield ratios for mineralised nutrients. So far as phytoplankton are concerned, it is best to conceive of a flexible nutrient element ratio for optimal growth, which depends on (a) relative values of subsistence quotas, and (b) internal storage capabilities. (Atomic) ratios based on subsistence quotas, and normalised to N = 16, are: N:Si:P:Fe = 16:(4–8–32):(0.3–0.6–1.0):(1–10 × 10−4) the silicon values applying, of course, only to diatoms. In northwest European seas, the natural ambient ratios of the inorganic macronutrients in winter seawater fall within the ranges . . . N:Si:P = 16:(6–20):(1–2.7) where the lack of Fe data reflects a present lack of adequate knowledge about the concentrations of assimilable iron in different parts of these seas. The main factors acting on these ratios have been discussed in Section 9.2; they include, in particular, denitrification and the nutrient element ratios in the adjacent ocean. And, of course, anthropogenic inputs of nitrogen by the atmospheric route and both N and P by way of rivers and direct discharges. Of these, the most widespread and substantial vehicle for anthropogenic nutrients is the discharge of rivers in which DAIN and total-P load is greatly enhanced, and dissolved silica is decreased compared to natural levels, with DIP influenced by suspended matter. Such discharges can perturb the ambient ratios of new nutrients outside the flexible ratio of algal growth given above and thus change the characteristic phytoplanktonic life-forms. However, the riverine discharge nutrient ratios impact significantly on the marine ratios only in nearshore regions where there is substantial freshwater content.

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Finally, a comparison of the natural ambient ratios with the algal growth ratios helps to confirm that new production (including that of diatoms) is expected to be nitrogen-limited. Box 1: Simplified cell-quota theory The theory is simplified from Droop (1983) and deals with population nutrient content or quota, Q (atom nutrient element (atom organic C)−1). Overall change includes biological change due to uptake of nutrient at rate u (at (at C)−1 d−1) and population growth at rate μ (d−1): dQ i, j ------------ = u i, j – μ i ⋅ Q i, j ± physical transport effects dt Sj u i, j = u maxi, j ⋅ ⎛ ------------------⎞ ⋅ f inhib ( Q i, j ) ⎝ k S + S j⎠

(1) (2)

i, j

where S is ambient nutrient concentration and the function finhib (Q) changes from 1 to 0 as the quota approaches a maximum Qmax. Suffix i gives the algal species and j the nutrient. The parameters are the maximum uptake rate umax and the seawater concentration kS at which uptake is half of this maximum. The term umax · (S/kS + S) can be derived by considering transport across the cell membrane at a finite number of sites. Because the ratio of cell surface to volume and biomass changes as (cell radius)2/3, larger cells have slower uptake relative to biomass than do smaller cells. Additionally, the diffusion of nutrient molecules across the viscous layer around a cell further diminishes the uptake ability of larger cells (Kiørboe, 1993). When ambient nutrient is scarce and quota is low compared with its maximum, uptake can be approximated by (3). When nutrient is plentiful, uptake is given by (4) Sj u i, j ≈ u maxi, j ⋅ ⎛ --------⎞ ⎝ kS ⎠

(3)

n ⎛ ⎞ ⎛ Q i, j ⎞ ⎟ ⎜ u i, j ≈ u maxi, j ⋅ 1 – ⎜ ---------------⎟ ⎜ ⎟ ⎝ Q max i, j⎠ ⎠ ⎝

(4)

i, j

The uptake inhibition term with an arbitrary power (n ≥ 1) is empirical (cf. Droop, 1983), and no account is taken here of uptake competition by several nutrients (cf. Flynn & Hipkin, 1999).

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Nutrient controlled growth is described by: k Qi, j⎞ ⎛ μ j = μ max j ⋅ ⎜ 1 – ρ i, j ⋅ -------Q i, j⎟⎠ ⎝

(5)

The parameters are maximum growth rate μmax, and minimum or subsistence quota kQ. The variable ρ is a luxury coefficient (ρ> 1 for non-limiting nutrient), empirically defined for a given species by: kQj Q j2 ρ j2 = --------1 ⋅ ------k Q j2 Q j1 (where j1 indicates the limiting nutrient and j2 the other). Light can also limit growth. Two classes of simplified solutions are useful. The first is relevant to a phytoplankton bloom, starting from high ambient nutrient concentration St = 0 and ending when this external nutrient has been exhausted and cellular concentration of limiting nutrient has fallen to the susbsistence quota. There is supposed to be no losses. In the case of a monospecific bloom, the final biomass is given by S j ,t = 0 [ +B j, t = 0 ⋅ Q j, t = 0 ] B maxj = ---------------------------------------------------------ρ j ⋅ k Qj

(6)

the square brackets enclosing initial algal nutrient. The final limiting nutrient is that for which ρ = 1. In a multi-species case, the succesful populations are those that get the largest share of the nutrient that proves to be limiting. They do this by (i) starting from a large initial population and (ii) being able to grow fast so as to keep Q < Qmax and hence allow sustained nutrient uptake. An approximate solution, for Sj1,t = 0 >> kSi,j is: μ .τ u maxi, j 1 Q i, j, t = 0 ⎞ ⎛⎛e t i . B maxi ≈ B i, t = 0 ⋅ ⎜ ---------- ⋅ k i, j1 ⋅ --------------⎞ + -----------------⎝ μi ⎠ k k Qi ⋅ j ⎟⎠ Qi, j1 ⎝

(7)

The parameter τi, gives the time at which the ambient nutrient concentration falls to that of the uptake half-saturation concentration for (limiting) nutrient j1. Population growth during this period is assumed to take place at a constant rate given by μi (d−1), more likely limited by light than by nutrient. The term ki takes account of the mean value of S/ks+S during this period, which is (by definition of the conditions) between 0.5 and 1. Mainly, however, it is influenced by the value of finhib (Qi). In effect, a species, or life-form, that grows with Q substantially less than Qmax, has a high value of k, which thus

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give its capability for luxury uptake. Thus, success amongst species in this yield-limited case is determined by the matrix of values of ki,j, umaxi,j and μi as well as those of kQi,j. The second class of solutions involves a steady state, which might occur in a laboratory chemostat, or in the upper waters of a tropical ocean where losses Li = μi are due to physical removal and grazing. In the steady state, ui,j = ρi,j ·μi · Qi,j. Given Sj << kSj, finhib (Q) ≈ 1. Growth rate is taken as half its maximum, so that Q = 2·ρ·kQ, and uj = μmax ·ρj ·kQj. For any particular species: u maxj 1 S j 1 1 u max j S j 1 ------------- ⋅ ------- ⋅ ------ ≈ μ max ≈ -------------2 ⋅ ------2- ⋅ -----k Qj1 k S j1 ρ j1 k Q j2 k Sj2 ρ j2 ρj ρ j1

(8)

S j k S j u max j k Qj S j2 k S j1 u max j1 k Qj2

⇒ρ* = ------2 ≈ ------1 ⋅ -------2 ⋅ -------------2 ⋅ --------1 where ρ* > 1 implies nutrient 1 limiting, and ρ* < 1 implies nutrient 2 limiting. S j 1 k Sj 1 u maxj 1 k Qj 2 Nutrient 1 is limiting when: ------ < ------- ⋅ ------------- ⋅ -------- . A solution for a limitS j2 k Sj2 u maxj2 k Qj1 ing nutrient and several species is: ℜ=

u maxi

- ⋅ B i⎞ ∑ ( u i ⋅ B i ) ≈ S ⋅ ∑ ⎛⎝ ---------⎠ kS i

i

(9)

i

where ℜ is the input of the nutrient (g-at m−3 d−1) due to vertical mixing and local remineralisation. In this uptake-limited case, the available nutrient supply is distributed amongst biota in proportion to their biomasses and their ratio of maximum uptake rate to half-saturation concentration. However, the steady state is unstable, as a small increase in biomass results in that species getting more of the available nutrient, leaving less for others: in principle the outcome is that all Bi except one become zero. Other species can survive if they are limited by different nutrients. Species limited by the same nutrient may also co-exist if grazing keeps their biomasses near a grazing threshold Bthr such that each population’s nutrient demand μi · Qi,j ·Bthr is much less than ℜ. Dedication Paul Tett wishes to dedicate this chapter to the memory of Mahlon G. Kelly, late of the University of Virginia, who introduced him to Biogeochemistry.

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Index

ABE 279 accretionary prisms 267 acid volatile sulphide 21, 71 acidifying gases 143 aerial roots 31 aerobic respiration 15 Alaska 145 Aleutian Trench 270 Alexandrium minutum 314 algal blooms 68 alien species 147 alkaline phosphatase 50 alveolata 304 Amazon river 243 ammonium 48 fluxes 17 anhydrite 242 anoxic conditions 165 reef pore-waters 53 waters 76 anthropogenic impacts on shallow reef 57 anthropogenic perturbations 276 anti-cyclonic circulation regime 139 anti-estuarine 91, 92, 93 aquaculture industry 68 ponds 1 Arabian Sea 157 Archaea 254, 303 Archaean cherts 260 Arctic 127 marginal seas 129 ocean 127 oscillation 139 shelves 133 aromatic hydrocarbons 268 assimilatory nitrate reduction 186 Asterionella 305 Aswan Dam 96 Atlantic waters 132 atmospheric carbon dioxide 224 atmospheric deposition 297 attenuation coefficient 172 Aureococcus 305

Aureoumbra 305 lagunensis 311 authigenic enrichments 85 autonomous vehicles 279 autotrophs 308 Axial Volcano 264 Bacillariophyceae 305 Bacillariophyta 305 back-arc spreading centres 241 back-reef 41 bacterial abundance 175 production 174 weathering 260 balance of organisms 294 ballast water 147 Baltic Sea 326, 340 Barents Sea 128, 136 barite 242, 269 basaltic glass 266 bathymodiolid mussels 240 Bay of Bengal 160 beam attenuation 171 Beggiatoa 268 benthic algae 8 carbon remineralisation 197 communities 195 nutrient pools 2 organisms 28 processes 195 benthic–pelagic coupling 137 benthos 137, 138, 228 Bering Strait 128 Bicoesida 305 bioamplification 144 biodiversity 138, 268 biofilms 255 biogeography 268 biomarkers 134 biomass 239 biome 345 biomineralisation 256 bioproduction 136

INDEX Biosphere 2 coral reef 45, 55 bioturbation 2, 20, 216, 218 Black Sea 116 black smokers 241 blooms 163 Bodo 304 bottom-up factors 312 brine 128 C:N ratio 319 C:N:P ratios 55 C:P ratio 46, 51 Cafeteria 305 Calanus helgolandicus 313 Canada 146 Canadian Archipelago 128 Basin 128 carbon cycle 263 cycling at vents 252 deposition 47 dioxide 185 dioxide flux to atmosphere 185 mineralization 12 reef 43, 44 carbonate–brucite chimneys 245 carbonates, microbial 274 carbonic acid 52 carotenoid 311 cell-quota theory 320 CEPEX 332 Ceratium 304, 310 Ceratium furca 312 Chaetoceros 305 chalcopyrite 242 Chatonella 305 chelation, organic 302 chemical tracers 76 chemocline 80 chemosynthetic 239 symbionts 267 chimney 254 Chlorococcales 305 Chlorophyceae 305 chlorophyll 107 Chlorophyta 305 Chloroxybacteria 304 choanoflagellates 305 chromium reducible sulphide 21 Chroococcales 304 Chrysochromulina 305, 309, 314 Chrysophyta 305

365

Chukchi Sea 128 Chytridomycota 305 chytrids 305 Ciliates 304 Ciliophora 304 cladoceran crustaceans 314 clams, Calyptogena 270 Solemya 270 climate change 138 feedback 278 models 141 shifts 139 variability 138 warming 141 coccolithophores 181 Coccoshpaerales 305 cod stocks 139 cold-condensation effect 144 cold-core eddy 96, 98 commercial fishing 145 shipping route 146 community production 46 respiratory quotients CRQ 15 competition experiments 337 competitive exclusion principle 307 contaminants 143 copepods 193, 313, 314 copper stress 314 co-precipitation 75 Corethron 169 Coscinodiscophyceae 305 Coscinodiscus 170, 310 coupled nitrification–denitrification 18 crabs 16 crude oil 268 cryptomonad 304, 311 Cryptomonas 304 Cryptophyta 304 CSR conceptual model 307 cyanobacteria 102, 304, 307, 311 Cyanophyceae 304 cycling of carbon 77 cyclonic circulation regime 139 CYPRIS station 300 Cyprus eddy 98 DAIN 294 decomposition constants 11 deep chlorophyll maximum (DCM) 91, 98, 100 denitrification 16, 50, 135, 165, 186, 299, 300 sedimentary 195

366 Desulfovibrio 272 detritus 7 diagenesis 69 diatom-copepod-fish trophic paradigm 313 diatoms 111, 170, 305 centric 307 pennate 307 tychopelagic 307 Dictyocha speculum 320 Dictyochophyceae 305 diffusive boundary layer 49, 54 DIN 166 Dinobryon 305 dinoflagellates 111, 304, 307, 314 Dinophyceae 304 Dinophysidae 304 Dinophysis 304 Dinophyta 304 Dissolved Available Inorganic Nitrogen 294 dissolved inorganic carbon 223 organic carbon 5, 47, 175 organic matter 4 organic nitrogen 5, 100, 298 organic phosphorus 100 organic pools 295 distribution coefficient 295 divergent plate boundaries 241 diversity 239, 293 DNA in vent fluids 246 doliolid salps 314 domoic acid 314 downwelling 166 DROOP model 320 East Greenland Current 131 ecological economics 144 ecosystem switch 187 eddies 160 Eel River Basin 273 El Niño–Southern Oscillation 140, 221, 222 electron transport system 187 element cycling 66 elemental sulphur 72 Emiliana 305 huxleyi 312 ENCORE experiments 57 environmental changes 138 enzymes, heat stable 251 eolian supply of dust 198 EPR 258

INDEX erosive margins 267 ERSEM 331 estuarine 296 Eubacteria 254, 303 Eucarya 303 eucaryotes 102 Euglenida 304 Eugleniphyceae 304 Euglenozoa 304 euphotic zone 102 Eurasian Basin 128 Eustigmatophyta 305 Eutreptiella 304 eutrophic equatorial abyss 214 event plumes 244 fast ice 132 fecal pellets 27, 184 feedback loops 138 fertilizer inputs from land 189 filaments 160 filtering effect 296 fish 140 processing 145 flood 75 floristic composition 307, 308, 310 flushing events 67 food web structure 179 foraminifera 181, 304 fore-reef 41 forest age 25 type 25 formaldehyde 51 fossil fuels 278 fossilised tube structures 259 Fragilariophyceae 305 Fram Srait 128 Framvaren Fjord 77 f-ratios 167 Fridtjof Nansen 140 galena 242 Gallionella 257 gas emission 185 hydrates 268 geological stability 276 global warming 141, 223 potential 278 Gonyaulacales 304 Gonyaulax 304

INDEX Gorda Ridge 259 greenhouse effect 132 Greenland 144, 146 Sea 131 groundwater discharge 297 Guaymas Basin 256, 274 Gulf of Mexico 270 Gymnodiniales 304 gyres 98 Gyrodinium 304 haptophytes 305, 307, 312 harbour seal 140 hazardous chemicals 143 heavy metals 143 Helgoland 329 herbivorous crabs 26 herring stocks 139 Heterokonts 305 Heterosigma 305 heterotrophic 239 bacteria 104, 174, 319 heterotrophic dinoflagellates 313, 319 heterotrophs 308 high carbon production 40 high nitrate low chlorophyll 345 high turbidity 5 Himmerfjord 337 hull fouling 147 hunters 145 Hydrate Ridge 273 hydraulic conductivity 53 hydrocarbon fluxes 281 hydrodynamics 3 hydrogen 264 oxidising bacteria 252 sulphide 239 hydrothermal fluids 242 heat flux 245 plumes 245, 252 hydrothermophilic microbes 246 ice algae 136 edge 131 fauna 136 formation 132 melting 132 ice-associated community 132 iceberg scour 138 Iheya Basin 258

367

impacts on coral reefs 58 incubation chambers 58 Indian Ocean 157 indigenous peoples 128, 144 Indo-Pacific reef flats 57 intermediate nepheloid layer 191 intraplate volcanism 244 iodate 51 iodine 51, 192 Irish Sea 296, 299, 300, 340 iron 192 associated with particles 302 colloidal 327 concentration, variations between oceans 301 deficiency 314 dissolved 301 enrichment 326 fertilisation 229 forms of 302 hydroxides 6 limitation 319, 345 limited 117 oxide muds 257 oxidisers 252 reduction 19 uptake 302 iron:phosphate ratio 323 iron-bound phosphate 24 iron-oxidising bacteria 261 Isochrysidales 305 isotope enrichments 78 biomarkers 281 fractionation 78 Japan Trench 267 Juan de Fuca Ridge 245 Kd 295 kelp forest 140 Kelvin waves 159 killer whale 140 K-selected 310 Labrador Sea 131, 139 laminated sediments 196 landslide 75 Lau basin 263 layer-silicates 258 leads 136 leaf-eating crabs 9

368 Leptocylindrus 305 Leptospirillum 257 Leptothrix 257 Levantine deep water 94 Levantine intermediate water 91 lifeforms 307 light limited 5 light-nutrient-mixing theory 309 Lingulodinium machaerophorum 310 lipid 136, 143 litter-bags 9 litter fall 7 Loch Creran 340 Loch Striven 340 LOICZ 299 Loihi Seamount 257 Lost City hydrothermal field 245 macrofauna 196 macrophyte production 8 magmatic degassing 241 management 146 manganese 192 nodules 218, 226 manganese-oxidising bacteria 252 mangrove crabs 7 creek 3 peats 23 mantle-derived methane 251 Manus basin 263 marcasite 242 marginal ice zone 135 marine mammals 145 mass flux 243 transfer relationships 55 MECHANISTIC models 325 Mediterranean outflow water 95 megafauna 228 Mendocino Fracture Zone 267 mercuric chloride 51 mercury pollution 73 Mersah Matru eddy 98 mesocosm experiments 332, 333 Mesodinium 304 mesophiles 251 mesoscale eddies 94 mesotrophic sub-equatorial abyss 217 mesozooplankton 177 metal mobility 240 metaliferous sediments 245 Metallogenium 257

INDEX metals, hydrothermal contribution to seawater 245 methane 132, 239 oxidation 83, 240, 272, 273 plumes 269 methane-producing bacteria 23 methanogenesis 239, 271 methanotrophs 272 methyl mercury 73 microbial layer 80 loop 184, 313 mineralisation 6 uptake 81 micro-organism taxonomy 303 microplankton 303 microzooplankton 175 grazing 180 Mid-Atlantic ridge 245 Middle Valley 256 mid-ocean-ridge 241 mining 277 mixed layer dynamics 172 layers 162 model, plankton 308 molecular phylogeny 240 MONOD theory 320 monsoon 2, 159 northeast 159 southwest 159 mother populations 277 mud volcanoes 267 myctophids 192 Myrionecta 304 myxotrophs 308 N fixing phytoplankton 118 N:P ratio 5, 53, 55, 311, 344 winter 301 N:Si ratio 332, 342 Nannochloropsis 305 native groups 145 natural capital 144 NeMO 279 net ecosystem production 299 neuston 111 new production 166 N-fixation 166 Nile 108 nitrate 48, 98 deficits 187 fluxes 18 uptake 183

INDEX nitrification 18, 50, 186 nitrite 48 reductase 319 nitrogen budget 298 fixation 50 oxidation states 294 reef 44 nitrogenase 326 Nitromonas 304 nitrous oxide 188 Nitzschia 170 Noctiluca 304 Noctilucales 304 nodule mining 226, 227 NOEWESP project 300 non-motile greens 307 nontronite 258 North Atlantic 165 net nitrogen flux 299 Oscillation 139 North Sea 296, 297, 312, 329 Project 297 Norway 146 Norwegian Sea 136 Nostocales 304 nutricline 97 nutrient concentrations 98 cycling 4 element ratios 319, 328 levels 4 limitation 116 loading of coral reefs 56 trap 164 uptake 46 nutrients 163 budget for North Sea 297 in coral reef waters 48 dissolved organic 312 sink 298 sorption to inorganic particles 295 oceanic biological pump 177 ochre 257 Ochromonas 305, 312 Ochromonodales 305 ODP 256 off-axis 243 oil exploration 277 Okinawa Trough 258 oligotrich ciliates 313 oligotrophic central gyre abyss 218 opal 181

ophiolites 243 Opsithokonts 305 Oregon margin 270 organic carbon accumulation 195 detritus 298 matter decomposition 16 matter preservation 85 nitrogen 48 Oscillatoriales 304 oxidation of H2S 82 oxides 244 oxygen conditions 67 consumption, sediment community 197 depletion 165 penetration 15 uptake 12, 13, 14, 15 oxygen-deficient zones 185 oxyhydroxides 242 Oxyrrhinaceae 304 P limited 116 Pacific Decadal Oscillation 221, 222 Pacific waters 132 pack ice 136 paradox of the plankton 307 Paraphysomonadaceae 305 particle aggregation 194 particle flux 180, 263 particulate organic matter 295 Pavlova 305 Pavlovales 305 PCB 69 pelagic heterotrophic community 6 tunicates 314 Pelagophyceae 305 Peridiniales 304 Peridinium trochoideum 310 peridotite 245 permanent ice pack 132 Persian Gulf 163 persistent organic contaminants 143 Peru margin 267, 270 petroleum 146, 268 pH changes 19 Phaeocystis 170, 305, 307, 314 Phagotrophic myxotrophs 312 phagotrophs 308 phosphate 98 authigenic 196 cycling 23 flux 24, 45

369

370 phosphorus enrichment 74 limited 91, 105 limiting 5 reef 44, 50 requirement 294 photoredox cycling 302 photosynthesis 45, 239 phycobilins 311 phycobiliprotein 304 physical disturbance 138 phytoplankton 8, 111 bio-optical properties 173 bloom 114, 173 community structure 167 composition 167 growth 180 growth and mortality 179 size distribution 167 pico-plankton 167 pigment composition 86 pigments 197 Pleuromamma indica 179 POC 163, 263 flux 211, 213, 220, 221, 225 polymetallic massive sulphides 277 polynyas 133 POM 295 pore water 72 pools 10 residence time 53 Porosira 169 Prasinophyceae 305 prawn farming 31 prey size 314 primary nitrite maximum 186 production 5, 105, 163, 299 productivity 97, 253 Prochlorococcus 169, 304, 311 prochlorophytes 102 Prorocentrales 304 Prorocentrum 304 micans 310 protected areas 278 Proteobacteria 304 Protoctista 303 Protoperidinium 304 prymeiophytes 104 Prymnesiales 305 Pseudomonas 304 Pseudo-nitzchia 305 Pseudo-nitzschia multiseries 314

INDEX pulp and paper industries 70 pyrite 71, 242 pyrite formation 82 Pyrocystales 304 Radiolaria 304 radionuclides 143 Rainbow MAR 263 Raphidophyta 305 rare earth element 78 reactive iron 72 recycled nutrients 329 Red Sea 163 red tide dinoflagellates 323 Redfield 315 Redfield ratio 6, 133, 293 flexible 347 redox 11, 213 potential 254 reactions of nitrogen 49 redoxcline 81 redox-sensitive elements 192 reducing conditions 189 reduction in iodate 192 reef crest 41 flat 41 morphology 41 zonation 41 reforestation 30 regime shifts 138 remineralisation 298 residence times 3 resilience 277 resource extraction 277 respiration, reef 45 respiratory quotient 45 reverse methanogenesis 83 Rhizosolenia 169, 310 Rhodes Gyre 95, 98 ridge flanks 243 River Po 115 river runoff 159 riverine inputs of nutrients 299 Rossby waves 159 r-selected 310 Russia 146 salps 184 salt tectonics 267 sapropels 92 saprotrophs 308

INDEX Sarcodina 304 Scandinavia 144 scavenging 73 Sea Cliff hydrothermal field 259 sea ice 127 lion 140 otter 140 urchin 140 seabirds 140 seafloor observatories 279 seagrass 29 leaves 11 seamounts 241 seasonal thermocline 100 variations 28 seawater-basalt interaction 242 SeaWiFS 113 sediment structure 12 trap 181 –water interface 68 sedimentation 163 rates 65 sediment–water interface 12, 26 Serial Endosymbiosis Theory 303 sesarmide crabs 26 sesarmine crab 2, 12 shear stress 3 shelf break 297 shelf-basin exchange 134 Shikmona Gyre 108 shrimp farming 7, 9, 23 Si/Ca ratio 181 Si:N:P ratio 311 Si:P ratio 311 siderophore 302, 325 silica 242 reef 51 requirement 319 silicate 98 requirement 294 Silicoflagellates 305 sill 67 size-fractionated chl a and primary production 168 Skagerrak 341 Skeletonema 305 slumping 65 Snake Pit 255 snowblower vents 246, 249 solar insolation 4 Southampton water 310

Southern Explorer 258 sphalerite 242 stable carbon isotopes 264 Stanton number 54 Straits of Sicily 94 Stramenophiles 305 stratification 173 stromatolites 275 Strombidium 304 suboxic sediments 76 suboxic zone 189 subsistence harvesting 145 quotas 321 subsurface biosphere 246 chlorophyll maximum 172 Sulfur Springs 260 sulphate reduction 15, 19, 22, 70, 81, 189, 240, 271 sulphate-reducers 251 sulphates 242 sulphide mounds 254 oxidation 240, 272 sulphides 242 sulphur oxidisers 249, 252 sulphur-rich floc 246 surface currents 159 suspended load 5 symbiosis 40 symbiotic 239 bacteria 48 Synechococcus 103, 169, 304, 311 Synura petersenii 320 Synurales 305 TAG 255 Thailand 1 Thalassiosira 305 thermohaline circulation 131 thermophilic organisms 246 Thioploca 196 thiosulphate 265 thiotrophs 272 tidal asymmetry 3 exchange 7 mixing fronts 309 Tintinnopsis 304 top-down trophic effects 312 Transpolar Drift 129 transport time in NW European shelf seas 301 Trichodesmium 166, 304

371

372 trophic behaviour 268 trophy 308 ultra-oligotrophic 91 ultraviolet-B radiation 127 underwater mining 277 United States 146 upwelling 160 urbanisation 9 varved sediments 196 vertical migration 179 vesicomyid clams 240 vestimentiferan worms 240 Vietnam 1 viral biomass 262 Viridioplantae 305 vitamin B12 312

INDEX volatiles 241 volcanic-hosted massive sulphide (VMS) 243 Volvocales 305 Wadden Sea 340 warm-core eddy 98 water, transport through reef framework 52 water-rock reactions 241 wave-induced mixing 53 West Spitsbergen Current 128 Western Mediterranean 93 wind stress 185 winter convection 166, 181 wurtzite 242 zooplankton 111, 264, 303 biomass 177, 179

Jan

Apr

Feb

May

Mar

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July

Oct

Aug

Nov

Sept

Dec

Color Plate 1 SeaWiFS 9-km monthly composite chlorophyll images for 2000; chlorophyll values in mg m−3 shown on the colour scale. SeaWiFS images courtesy of NASA SeaWiFS Project and Orbimage Inc. (see Fig. 4.7).

Chl a (mg/m3)

Temperature (°C)

25 20 15 10 5

Latitude (N)

25 20 15 10 5 25 20 15 10 5 50

0.0

55

0.4

60

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65

70

1.5

75 50 55 Longitude (E) 4.0

18

22

60

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26

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75

34

Color Plate 2 Monthly climatologies of remotely sensed surface chlorophyll a (a–c) and SST (d–f) for February (a, d), May (b, e) and August (c, f). Chlorophyll and SST climatologies were created by Jerry Wiggert (University of Maryland, College Park) and Bob Evans (RSMAS, University of Miami) using the SeaWiFS (Sea-viewing Wide Field-of-view Sensor) and MODIS (Moderate-resolution Imaging Spectroradiometer) data, respectively. Note that some features of circulation (e.g. the great whirl) can be readily identified in both sets of data (see Fig. 6.2).

Color Plate 3 Distributions of chlorophyll fluorescence (panels on the left) and POC (panels on the right) in the upper 150 m along the US JGOFS southern transect during different seasons. These properties were derived from in situ light transmission and fluorescence measurements. The dotted and solid lines indicate the mixed layer depth (identified by an increase in σθ by 0.03 relative to the surface) and the 0.5 μM NO3− contour, respectively. Reproduced from Gundersen et al. (1998) with permission from Elsevier Science (see Fig. 6.8).

Color Plate 4 Acoustic backscatter for 24 hours around 19° N, 67° E indicating diurnal migration of organisms (mostly myctophids). Reproduced from Morrison et al. (1999) (see Fig. 6.18).

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