ENCYCLOPEDIA OF
OCEAN SCIENCES SECOND EDITION
Editor-in-chief
JOHN H. STEELE
Editors
STEVE A. THORPE KARL K. TUREKIAN
Boston • Heidelberg • London • New York • Oxford Paris • San Diego • San Francisco • Singapore • Sydney • Tokyo Academic Press is an imprint of Elsevier
(c) 2011 Elsevier Inc. All Rights Reserved.
ENCYCLOPEDIA OF
OCEAN SCIENCES SECOND EDITION
(c) 2011 Elsevier Inc. All Rights Reserved.
Subject Area Volumes from the Second Edition Climate & Oceans edited by Karl K. Turekian Elements of Physical Oceanography edited by Steve A. Thorpe Marine Biology edited by John H. Steele Marine Chemistry & Geochemistry edited by Karl K. Turekian Marine Ecological Processes edited by John H. Steele Marine Geology & Geophysics edited by Karl K. Turekian Marine Policy & Economics guest edited by Porter Hoagland, Marine Policy Center, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts Measurement Techniques, Sensors & Platforms edited by Steve A. Thorpe Ocean Currents edited by Steve A. Thorpe The Coastal Ocean edited by Karl K. Turekian The Upper Ocean edited by Steve A. Thorpe
(c) 2011 Elsevier Inc. All Rights Reserved.
ENCYCLOPEDIA OF
OCEAN SCIENCES SECOND EDITION Volume 3: G - M Editor-in-chief
JOHN H. STEELE
Editors
STEVE A. THORPE KARL K. TUREKIAN
Boston • Heidelberg • London • New York • Oxford Paris • San Diego • San Francisco • Singapore • Sydney • Tokyo Academic Press is an imprint of Elsevier
(c) 2011 Elsevier Inc. All Rights Reserved.
Academic Press is an imprint of Elsevier 32 Jamestown Road, London NW1 7BY, UK 30 Corporate Drive, Suite 400, Burlington, MA 01803, USA 525 B Street, Suite 1900, San Diego, CA 92101-4495, USA Copyright ^ 2009 Elsevier Ltd. All rights reserved
The following articles are US government works in the public domain and are not subject to copyright: Fish Predation and Mortality; International Organizations; Large Marine Ecosystems; Ocean Circulation: Meridional Overturning Circulation; Salt Marsh Vegetation; Satellite Passive-Microwave Measurements of Sea Ice; Satellite Oceanography, History and Introductory Concepts; Satellite Remote Sensing: Ocean Color; Science of Ocean Climate Models; Wind- and Buoyancy-Forced Upper Ocean. Fish Migration, Horizontal Crown Copyright 2001 Turbulence Sensors Canadian Crown Copyright 2001 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 without the prior written permission of the publisher
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ISBN: 978-0-12-375044-0
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Editors
Editor-in-chief John H. Steele Marine Policy Center, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA
Editors Steve A. Thorpe National Oceanography Centre, University of Southampton Southampton, UK School of Ocean Sciences, University of Bangor, Menai Bridge, Anglesey, UK Karl K. Turekian Yale University, Department of Geology and Geophysics, New Haven, Connecticut, USA
(c) 2011 Elsevier Inc. All Rights Reserved.
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Editorial Advisory Board John H. S. Blaxter Scottish Association for Marine Science Dunstaffnage Marine Laboratory Oban Argyll, UK Quentin Bone The Marine Biological Association of the United Kingdom Plymouth, UK Kenneth H. Brink Woods Hole Oceanographic Institution Woods Hole MA, USA Harry L. Bryden School of Ocean and Earth Science James Rennell Division University of Southampton Empress Dock Southampton, UK Robert Clark University of Newcastle upon Tyne Marine Sciences and Coastal Management Newcastle upon Tyne, UK J. Kirk Cochran State University of New York at Stony Brook Marine Sciences Research Center Stony Brook NY, USA Jeremy S. Collie Coastal Institute Graduate School of Oceanography University of Rhode Island South Ferry Road Narragansett RI, USA
Paul G. Falkowski Departments of Geological Sciences & Marine & Coastal Sciences Institute of Marine & Coastal Sciences School of Environmental & Biological Sciences Rutgers University New Brunswick NJ, USA Mike Fashamw Southampton Oceanography Centre University of Southampton Southampton UK John G. Field MArine REsearch (MA-RE) Institute University of Cape Town Rondebosch South Africa Michael Fogarty NOAA, National Marine Fisheries Service Woods Hole MA, USA Wilford D. Gardner Department of Oceanography Texas A&M University College Station TX, USA Ann Gargett Old Dominion University Center for Coastal Physical Oceanography Crittenton Hall Norfolk VA, USA
Peter J. Cook Australian Petroleum Cooperative Research Centre Canberra, Australia
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Robert A. Duce Departments of Oceanography and Atmospheric Sciences Texas A&M University College Station TX, USA
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deceased
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Editorial Advisory Board
Christopher Garrett University of Victoria Department of Physics Victoria British Columbia, Canada
Lindsay Lairdw Aberdeen University Zoology Department Aberdeen UK
W. John Gould Southampton Oceanography Centre University of Southampton Southampton UK
Peter S. Liss University of East Anglia School of Environmental Sciences Norwich, UK
John S. Grayw Institute of Marine Biology and Limnology University of Oslo Blindern Oslo, Norway
Ken Macdonald University of California Department of Geological Sciences Santa Barbara CA, USA
Gwyn Griffiths Southampton Oceanography Centre University of Southampton Southampton UK
Dennis McGillicuddy Woods Hole Oceanographic Institution Woods Hole MA, USA Alasdair McIntyre University of Aberdeen Department of Zoology Aberdeen UK
Stephen J. Hall World Fish Center Penang Malaysia Roger Harris Plymouth Marine Laboratory West Hoe Plymouth, UK Porter Hoagland Woods Hole Oceanographic Institution Woods Hole MA, USA George L. Hunt Jr. University of California, Irvine Department of Ecology and Evolutionary Biology Irvine CA, USA William J. Jenkins Woods Hole Oceanographic Institution Woods Hole MA, USA
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deceased
W. Kendall Melville Scripps Institution of Oceanography UC San Diego La Jolla CA, USA John Milliman College of William and Mary School of Marine Sciences Gloucester Point VA, USA James N. Moum College of Oceanic and Atmospheric Sciences Oregon State University Corvallis OR, USA Michael M. Mullinw Scripps Institution of Oceanography Marine Life Research Group University of California San Diego La Jolla CA, USA
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Editorial Advisory Board
Yoshiyuki Nozakiw University of Tokyo The Ocean Research Institute Nakano-ku Tokyo Japan
Ellen Thomas Yale University Department of Geology and Geophysics New Haven CT, USA
John Orcutt Scripps Institution of Oceanography Institute of Geophysics and Planetary Physics La Jolla CA, USA Richard F. Pittenger Woods Hole Oceanographic Institution Woods Hole MA, USA Gerold Siedler Universita¨t Kiel Institut fua¨r Meereskunde Kiel Germany
Peter L. Tyack Woods Hole Oceanographic Institution Woods Hole MA, USA Bruce A. Warren Woods Hole Oceanographic Institution Woods Hole MA, USA Wilford F. Weeks University of Alaska Fairbanks Department of Geology and Geophysics Fairbanks AK, USA
Robert C. Spindel University of Washington Applied Physics Laboratory Seattle WA, USA
Robert A. Weller Woods Hole Oceanographic Institution Woods Hole MA, USA
Colin P. Summerhayes Scientific Committee on Antarctic Research (SCAR) Scott Polar Institute Cambridge, UK
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Stewart Turner Australian National University Research School of Earth Sciences Canberra Australia
James A. Yoder Woods Hole Oceanographic Institution Woods Hole MA, USA
deceased
(c) 2011 Elsevier Inc. All Rights Reserved.
Preface to Second Print Edition The first edition of the Encyclopedia of Ocean Sciences, published in print form in 2001, has proven to be a valuable asset for the marine science community – and more generally. The continuing rapid increase in electronic access to academic material led us initially to publish the second edition electronically. We have now added this print version of the second edition because of a demonstrated need for such a product. The encyclopedia can now be accessed in print or electronic format according to the preferences and needs of individuals and institutions. In this edition there are 54 new articles, 67 revisions of previous articles, and a completely revised and improved index. We are grateful to the members of the Editorial Advisory Board, nearly all of whom have stayed with us during the lengthy process of going electronic. The transition from Academic Press to Elsevier occurred between the two editions. We thank Dr. Debbie Tranter of Elsevier for her efforts to see this edition through its final stages.
Preface to First Edition In 1942, a monumental volume was published on The Oceans by H. U. Sverdrup, M. W. Johnson, and R. H. Fleming. It was comprehensive and covered the knowledge at that time of the scientific study of the oceans. This seminal book helped to initiate the tremendous burgeoning of marine research that occurred during the following decades. The Encyclopedia of Ocean Sciences aims to embody the great growth of knowledge in a major new reference work. There have been remarkable new approaches to the study of the oceans that blur the distinctions between the physical, chemical, biological, and geological disciplines. New theories and technologies have expanded our knowledge of ocean processes. For example, plate tectonics has revolutionized our view not only of the geology and geophysics of the seafloor but also of ocean chemistry and biology. Satellite remote sensing provides a global vision as well as detailed understanding of the close coupling of ocean physics and biology at local and regional scales. Exploration, fishing, warfare, and the impact of storms have driven the past study of the seas, but we now have a great public awareness of and concern with broader social and economic issues affecting the oceans. For this reason, we have invited articles explicitly on marine policy and environmental topics, as well as encouraged authors to address these aspects of their particular subjects. We believe the encyclopedia should be of use to those involved with policy and management as well as to students and researchers. Over 400 scientists have contributed to this description of what we now know about the oceans. They are distinguished researchers who have generously shared their knowledge of this ever-growing body of science. We are extremely grateful to all these authors, whose ability to write concisely on complex subjects has generated a perspective on our science that we, as editors, believe will enhance the appreciation of the oceans, their uses, and the research ahead. It has been a major challenge for the members of the Editorial Advisory Board to cover such a heterogeneous subject. Their knowledge of the diverse areas of research has guaranteed comprehensive coverage of the ocean sciences. The Board contributed significantly by suggesting topics, persuading authors to contribute, and reviewing drafts. Many of them wrote Overviews that give broad descriptions of major parts of the ocean sciences. Clearly, it was the dedicated involvement of the Editorial Advisory Board that made this venture successful. Such a massive enterprise as a multivolume encyclopedia would not be possible without the long-term commitment of the staff of the Major Reference Works team at Academic Press. In particular, we are very grateful for the consistent support of our Senior Developmental Editor, Colin McNeil, who has worked so well with us throughout the whole process. Also, we are very pleased that new technology permits enhanced search and retrieval through the Internet. We believe this will make the encyclopedia much more accessible to individual researchers and students.
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Preface to Second Print Edition
In Memoriam During the creation of the Encyclopedia of Ocean Sciences and also in several cases prior to the publication of the electronic Second Edition, several Associate Editors or designated Associate Editors died. We specifically acknowledge their role in making this work an effective publication. They are Mike Fasham, John S. Gray, Lindsay Laird, Michael Mullin and Yoshiyuki Nozaki. J. H. Steele, S. A. Thorpe, and K. K. Turekian Editors
(c) 2011 Elsevier Inc. All Rights Reserved.
Guide to Use of the Encyclopedia
Introductory Points In devising the vision and structure for the Encyclopedia, the Editors have striven to unite and interrelate all current knowledge that can be designated ‘‘Ocean Sciences’’. To aid users of the Encyclopedia, this new reference work offers intuitive searching and extensive cross-linking of content. These features are explained in more detail below.
Structure of the Encyclopedia The material in the Encyclopedia is arranged as a series of articles in alphabetical order. To help you realize the full potential of the material in the Encyclopedia we have provided three features to help you find the topic of your choice.
1. Contents Lists Your first point of reference will probably be the contents list. The contents list appearing in each volume will provide you with the page number of the article. Alternatively you may choose to browse through a volume using the alphabetical order of the articles as your guide. To assist you in identifying your location within the Encyclopedia a running headline indicates the current article.
2. Cross References All of the articles in the encyclopedia have heen extensively cross referenced. The cross references, which appear at the end of each article, have heen provided at three levels: i. To indicate if a topic is discussed in greater detail elsewhere.
ACOUSTICS, ARCTIC See also: Acoustics in Marine Sediments. Acoustic Noise. Acoustics, Shallow Water. Arctic Ocean Circulation. Bioacoustics. Ice–ocean interaction. Nepheloid Layers. North Atlantic Oscillation (NAO). Ocean Circulation: Meridional Overturning Circulation. Platforms: Autonomous Underwater Vehicles. Satellite Passive-Microwave Measurements of Sea Ice. Sea Ice. Sea Ice: Overview. Seals. Seismic Structure. Tomography. Under-Ice Boundary Layer. Water Types and Water Masses.
ii. To draw the reader’s attention to parallel discussions in other articles. ACOUSTICS, ARCTIC See also: Acoustics in Marine Sediments. Acoustic Noise. Acoustics, Shallow Water. Arctic Ocean Circulation. Bioacoustics. Ice–ocean interaction. Nepheloid Layers. North Atlantic Oscillation (NAO). Ocean Circulation: Meridional Overturning Circulation. Platforms: Autonomous Underwater Vehicles. Satellite Passive-Microwave Measurements of Sea Ice. Sea Ice. Sea Ice: Overview. Seals. Seismic Structure. Tomography. Under-Ice Boundary Layer. Water Types and Water Masses.
(c) 2011 Elsevier Inc. All Rights Reserved.
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Guide to Use of the Encyclopedia
iii. To indicate material that broadens the discussion.
ACOUSTICS, ARCTIC See also: Acoustics in Marine Sediments. Acoustic Noise. Acoustics, Shallow Water. Arctic Ocean Circulation. Bioacoustics. Ice–ocean interaction. Nepheloid Layers. North Atlantic Oscillation (NAO). Ocean Circulation: Meridional Overturning Circulation. Platforms: Autonomous Underwater Vehicles. Satellite Passive-Microwave Measurements of Sea Ice. Sea Ice. Sea Ice: Overview. Seals. Seismic Structure. Tomography. Under-Ice Boundary Layer. Water Types and Water Masses.
3. Index The index will provide you with the volume and page number where the material is to be located, and the index entries differentiate between material that is a whole article, is part of an article or is data presented in a table or figure. On the opening page of the index detailed notes are provided.
4. Appendices In addition to the articles that form the main body of the encyclopedia, there are a number of appendices which provide bathymetric charts and lists of data used throughout the encyclopedia. The appendices are located in volume 6, before the index.
5. Contributors A full list of contributors appears at the beginning of volume 1.
(c) 2011 Elsevier Inc. All Rights Reserved.
Contributors Volume 1 E E Adams
N Caputi
Massachusetts Institute of Technology, Cambridge, MA, USA
Fisheries WA Research Division, North Beach, WA, Australia
T Akal
C A Carlson
NATO SACLANT Undersea Research Centre, La Spezia, Italy
University of California, Santa Barbara, CA, USA H Chamley
R Arimoto New Mexico State University, Carlsbad, NM, USA
Universite´ de Lille 1, Villeneuve d’Ascq, France R Chester
J L Bannister The Western Australian Museum, Perth, Western Australia
Liverpool University, Liverpool, Merseyside, UK V Christensen University of British Columbia, Vancouver, BC, Canada
E D Barton University of Wales, Bangor, UK
J W Dacey
N R Bates Bermuda Biological Station for Research, St George’s, Bermuda, USA
Woods Hole Oceanographic Institution, Woods Hole, MA, USA R A Duce
A Beckmann
Texas A&M University, College Station, TX, USA
Alfred-Wegener-Institut fu¨r Polar und Meeresforschung, Bremerhaven, Germany
H W Ducklow
P S Bell
The College of William and Mary, Gloucester Point, VA, USA
Proudman Oceanographic Laboratory, Liverpool, UK I Dyer G Birnbaum
Marblehead, MA, USA
Alfred-Wegener-Institut fu¨r Polar und Meeresforschung, Bremerhaven, Germany
D W Dyrssen Gothenburg University, Go¨teborg, Sweden
B O Blanton The University of North Carolina at Chapel Hill, Chapel Hill, NC, USA E A Boyle Massachusetts Institute of Technology, Cambridge, MA, USA
S M Evans Newcastle University, Newcastle, UK I Everson Anglia Ruskin University, Cambridge, UK
P Boyle
J W Farrington
University of Aberdeen, Aberdeen, UK
Woods Hole Oceanographic Institution, MA, USA
D M Bush
M Fieux
State University of West Georgia, Carrollton, GA, USA
Universite´ Pierre et Marie Curie, Paris, France
K Caldeira
R A Fine
Stanford University, Stanford, CA, USA
University of Miami, Miami, FL, USA
(c) 2011 Elsevier Inc. All Rights Reserved.
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Contributors
K G Foote Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA L Franc¸ois University of Lie`ge, Lie`ge, Belgium M A M Friedrichs Old Dominion University, Norfolk, VA, USA T Gaston National Wildlife Research Centre, Quebec, Canada J Gemmrich University of Victoria, Victoria, BC, Canada Y Godde´ris University of Lie`ge, Lie`ge, Belgium D R Godschalk University of North Carolina, Chapel Hill, NC, USA A J Gooday Southampton Oceanography Centre, Southampton, UK A L Gordon Lamont-Doherty Earth Observatory of Columbia University, Palisades, NY, USA D A Hansell University of Miami, Miami FL, USA L W Harding Jr University of Maryland, College Park, MD, USA R Harris Plymouth Marine Laboratory, Plymouth, UK P J Herring Southampton Oceanography Centre, Southampton, UK B M Hickey University of Washington, Seattle, WA, USA M A Hixon Oregon State University, Corvallis, OR, USA E E Hofmann Old Dominion University, Norfolk, VA, USA S Honjo Woods Hole Oceanographic Institution, Woods Hole, MA, USA D J Howell Newcastle University, Newcastle, UK J M Huthnance CCMS Proudman Oceanographic Laboratory, Wirral, UK B Ja¨hne University of Heidelberg, Heidelberg, Germany F B Jensen SACLANT Undersea Research Centre, La Spezia, Italy A John Sir Alister Hardy Foundation for Ocean Science, Plymouth, UK
C D Jones University of Washington, Seattle, WA, USA P F Kingston Heriot-Watt University, Edinburgh, UK W Krauss Institut fu¨r Meereskunde an der Universita¨t Kiel, Kiel, Germany W A Kuperman Scripps Institution of Oceanography, University of California, San Diego, CA, USA D Lal Scripps Institute of Oceanography, University of California San Diego, La Jolla, CA, USA C S Law Plymouth Marine Laboratory, The Hoe, Plymouth, UK W J Lindberg University of Florida, Gainesville, FL, USA J R E Lutjeharms University of Cape Town, Rondebosch, South Africa P Malanotte-Rizzoli Massachusetts Institute of Technology, Cambridge, MA, USA W R Martin Woods Hole Oceanographic Institution, Woods Hole, MA, USA R P Matano Oregon State University, Corvallis, OR, USA J W McManus University of Miami, Miami, FL, USA G M McMurtry University of Hawaii at Manoa, Honolulu, HI, USA R Melville-Smith Fisheries WA Research Division, North Beach, WA, Australia P N Mikhalevsky Science Applications International Corporation, McLean, VA, USA W D Miller University of Maryland, College Park, MD, USA D Monahan University of New Hampshire, Durham, NH, USA J C Moore University of California at Santa Cruz, Santa Cruz, CA, USA A Morel Universite´ Pierre et Marie Curie, Villefranche-sur-Mer, France R Narayanaswamy The University of Manchester, Manchester, UK
(c) 2011 Elsevier Inc. All Rights Reserved.
Contributors W J Neal Grand Valley State University, Allendale, MI, USA D Pauly University of British Columbia, Vancouver, BC, Canada J W Penn Fisheries WA Research Division, North Beach, WA, Australia L C Peterson University of Miami, Miami, FL, USA S G Philander Princeton University, Princeton, NJ, USA N J Pilcher Universiti Malaysia Sarawak, Sarawak, Malaysia O H Pilkey Duke University, Durham, NC, USA
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D H Shull Western Washington University, Bellingham, WA, USA D K Steinberg College of William and Mary, Gloucester Pt, VA, USA L Stramma University of Kiel, Kiel, Germany R N Swift NASA Goddard Space Flight Center, Wallops Island, VA, USA T Takahashi Lamont Doherty Earth Observatory, Columbia University, Palisades, NY, USA P D Thorne Proudman Oceanographic Laboratory, Liverpool, UK
A R Piola Universidad de Buenos Aires, Buenos Aires, Argentina J M Prospero University of Miami, Miami, FL, USA S Rahmstorf Potsdam Institute for Climate Impact Research, Potsdam, Germany P C Reid SAHFOS, Plymouth, UK G Reverdin LEGOS, Toulouse Cedex, France S R Rintoul CSIRO Antarctic Climate and Ecosystems Cooperative Research Centre, Hobart, TAS, Australia J M Roberts Scottish Association for Marine Science, Oban, UK P A Rona Rutgers University, New Brunswick, NJ, USA T C Royer Old Dominion University, Norfolk, VA, USA B Rudels Finnish Institute of Marine Research, Helsinki, Finland
P L Tyack Woods Hole Oceanographic Institution, Woods Hole, USA T Tyrrell National Oceanography Centre, Southampton, UK F E Werner The University of North Carolina at Chapel Hill, Chapel Hill, NC, USA E A Widder Harbor Branch Oceanographic Institution, Fort Pierce, FL, USA D J Wildish Fisheries and Oceans Canada, St. Andrews, NB, Canada A J Williams, III Woods Hole Oceanographic Institution, Woods Hole, MA, USA D K Woolf Southampton Oceanography Centre, Southampton, UK
W Seaman University of Florida, Gainesville, FL, USA
C W Wright NASA Goddard Space Flight Center, Wallops Island, VA, USA
F Sevilla, III, University of Santo Tomas, Manila,The Philippines
J D Wright Rutgers University, Piscataway, NJ, USA
L V Shannon University of Cape Town, Cape Town, South Africa
J R Young The Natural History Museum, London, UK
G I Shapiro University of Plymouth, Plymouth, UK
H J Zemmelink University of Groningen, Haren, The Netherlands
A D Short University of Sydney, Sydney, Australia
W Zenk Universita¨t Kiel, Kiel, Germany
(c) 2011 Elsevier Inc. All Rights Reserved.
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Contributors
Volume 2 G P Arnold Centre for Environment, Fisheries & Aquaculture Science, Suffolk, UK
K Dyer University of Plymouth, Plymouth, UK
K M Bailey Alaska Fisheries Science Center, Seattle, WA, USA
M Elliott Institute of Estuarine and Coastal Studies, University of Hull, Hull, UK
J G Baldauf Texas A&M University, College Station, TX, USA
D M Farmer Institute of Ocean Sciences, Sidney, BC, Canada
J Bascompte CSIC, Seville, Spain
A V Fedorov Yale University, New Haven, CT, USA
A Belgrano Institute of Marine Research, Lysekil, Sweden
M J Fogarty Northeast Fisheries Science Center, National Marine Fisheries Service, Woods Hole, MA, USA
O A Bergstad Institute of Marine Research, Flødevigen His, Norway J H S Blaxter Scottish Association for Marine Science, Argyll, UK
R Fonteyne Agricultural Research Centre, Ghent, Oostende, Belgium
Q Bone The Marine Biological Association of the United Kingdom, Plymouth, UK
D J Fornari Woods Hole Oceanographic Institution, Woods Hole, USA
I Boyd University of St. Andrews, St. Andrews, UK
A E Gargett Old Dominion University, Norfolk, VA, USA
K M Brander DTU Aqua, Charlottenlund, Denmark and International Council for the Exploration of the Sea (ICES), Copenhagen, Denmark
C H Gibson University of California, San Diego, La Jolla, CA, USA
J N Brown Yale University, New Haven, CT, USA T K Chereskin University of California San Diego, La Jolla, CA, USA J S Collie Danish Institute for Fisheries Research, Charlottenlund, Denmark and University of Rhode Island, Narragansett, RI, USA G Cresswell CSIRO Marine Research, Tasmania, Australia
J D M Gordon Scottish Association for Marine Science, Argyll, UK J F Grassle Rutgers University, New Brunswick, New Jersey, USA S J Hall Flinders University, Adelaide, SA, Australia N Hanson University of St. Andrews, St. Andrews, UK P J B Hart University of Leicester, Leicester, UK
J Davenport University College Cork, Cork, Ireland
K R Helfrich Woods Hole Oceanographic Institution, Woods Hole, MA, USA
R H Douglas City University, London, UK
D M Higgs University of Windsor, Windsor, ON, Canada
S Draxler Karl-Franzens-Universita¨t Graz, Graz, Austria
N G Hogg Woods Hole Oceanographic Institution, Woods Hole, MA, USA
J T Duffy-Anderson Alaska Fisheries Science Center, Seattle, WA, USA J A Dunne Santa Fe Institute, Santa Fe, NM, USA and Pacific Ecoinformatics and Computational Ecology Lab, Berkely, CA, USA
E D Houde University of Maryland, Solomons, MD, USA V N de Jonge Department of Marine Biology, Groningen University, Haren, The Netherlands
(c) 2011 Elsevier Inc. All Rights Reserved.
Contributors K Katsaros Atlantic Oceanographic and Meteorological Laboratory, NOAA, Miami, FL, USA J M Klymak University of Victoria, Victoria, BC, Canada M Kucera Eberhard Karls Universita¨t Tu¨bingen, Tu¨bingen, Germany R S Lampitt University of Southampton, Southampton, UK J R N Lazier Bedford Institute of Oceanography, NS, Canada J R Ledwell Woods Hole Oceanographic Institution, Woods Hole, MA, USA P F J Lermusiaux Harvard University, Cambridge, MA, USA M E Lippitsch Karl-Franzens-Universita¨t Graz, Graz, Austria
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T J Pitcher University of British Columbia, Vancouver, Canada A N Popper University of Maryland, College Park, MD, USA J F Price Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA R D Prien Southampton Oceanography Centre, Southampton, UK A-L Reysenbach Portland State University, Portland, OR, USA P L Richardson Woods Hole Oceanographic Institution, Woods Hole, MA, USA A R Robinson Harvard University, Cambridge, MA, USA M D J Sayer Dunstaffnage Marine Laboratory, Oban, Argyll, UK
B J McCay Rutgers University, New Brunswick, NJ, USA
R W Schmitt Woods Hole Oceanographic Institution, Woods Hole, MA, USA
J D McCleave University of Maine, Orono, ME, USA
J Scott DERA Winfrith, Dorchester, Dorset, UK
D Minchin Marine Organism Investigations, Killaloe, Republic of Ireland
M P Sissenwine Northeast Fisheries Science Center, Woods Hole, MA, USA
C M Moore University of Essex, Colchester, UK K Moran University of Rhode Island, Narragansett, RI, USA G R Munro University of British Columbia, Vancouver, BC, Canada J D Nash Oregon State University, Corvallis, Oregon, OR, USA A C Naveira Garabato University of Southampton, Southampton, UK
T P Smith Northeast Fisheries Science Center, Woods Hole, MA, USA P V R Snelgrove Memorial University of Newfoundland, St John’s, NL, Canada M A Spall Woods Hole Oceanographic Institution, Woods Hole, MA, USA A Stigebrandt University of Gothenburg, Gothenburg, Sweden D A V Stow University of Southampton, Southampton, UK
J D Neilson Department of Fisheries and Oceans, New Brunswick, Canada
D J Suggett University of Essex, Colchester, UK
Y Nozakiw University of Tokyo, Tokyo, Japan
U R Sumaila University of British Columbia, Vancouver, BC, Canada
R I Perry Department of Fisheries and Oceans, British Columbia, Canada S G Philander Princeton University, Princeton, NJ, USA w
Deceased.
K S Tande Norwegian College of Fishery Science, Tromsø, Norway S A Thorpe National Oceanography Centre, Southampton, UK R S J Tol Economic and Social Research Institute, Dublin, Republic of Ireland
(c) 2011 Elsevier Inc. All Rights Reserved.
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Contributors
K E Trenberth National Center for Atmospheric Research, Boulder, CO, USA J J Videler Groningen University, Haren, The Netherlands
R S Wells Chicago Zoological Society, Sarasota, FL, USA D C Wilson Institute for Fisheries Management and Coastal Community Development, Hirtshals, Denmark
Volume 3 S Ali Plymouth Marine Laboratory, Plymouth, UK
K H Coale Moss Landing Marine Laboratories, CA, USA
J T Andrews University of Colorado, Boulder, CO, USA
M F Coffins University of Texas at Austin, Austin, TX, USA
M A de Angelis Humboldt State University, Arcata, CA, USA
P J Corkeron James Cook University, Townsville, Australia
A J Arp Romberg Tiburon Center for Environment Studies, Tiburon, CA, USA
B C Coull University of South Carolina, Columbia, SC, USA
T Askew Harbor Branch Oceanographic Institute, Ft Pierce, FL, USA
R Cowen University of Miami, Miami, FL, USA
R D Ballard Institute for Exploration, Mystic, CT, USA
G Cresswell CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia
G Barnabe´ Universite´ de Montpellier II, France
D S Cronan Royal School of Mines, London, UK
R S K Barnes University of Cambridge, Cambridge, UK
J Csirke Food and Agriculture Organization of the United Nations, Rome, Italy
E D Barton University of Wales, Bangor, Menai Bridge, Anglesey, UK
G A Cutter Old Dominion University, Norfolk, VA, USA
D Bhattacharya University of Iowa, Iowa City, IA, USA
D J DeMaster North Carolina State University, Raleigh, NC, USA
F von Blanckenburg Universita¨t Bern, Bern, Switzerland
T D Dickey University of California, Santa Barbara, CA, USA
D R Bohnenstiehl North Carolina State University, Raleigh, NC, USA
D Diemand Coriolis, Shoreham, VT, USA
H L Bryden University of Southampton, Southampton, UK J Burger Rutgers University, Piscataway, NJ, USA S M Carbotte Lamont-Doherty Earth Observatory of Columbia University, Palisades, NY, USA G T Chandler University of South Carolina, Columbia, SC, USA M A Charette Woods Hole Oceanographic Institution, Woods Hole, MA, USA
C S M Doake British Antarctic Survey, Cambridge, UK C M Domingues CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia C J Donlon Space Applications Institute, Ispra, Italy F Doumenge Muse´e Oce´anographique de Monaco, Monaco R A Dunn University of Hawaii at Manoa, Honolulu, HI, USA
(c) 2011 Elsevier Inc. All Rights Reserved.
Contributors R P Dziak Oregon State University/National Oceanic and Atmospheric Administration, Hatfield Marine Science Center, Newport, OR, USA O Eldholm University of Oslo, Oslo, Norway
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S K Hooker University of St. Andrews, St. Andrews, UK H Hotta Japan Marine Science & Technology Center, Japan G R Ierley University of California San Diego, La Jolla, CA, USA
A E Ellis Marine Laboratory, Aberdeen, Scotland, UK C R Engle University of Arkansas at Pine Bluff, Pine Bluff, AR, USA C C Eriksen University of Washington, Seattle, WA, USA V Ettwein University College London, London, UK S Farrow Carnegie Mellon University, Pittsburgh, PA, USA M Fieux Universite´ Pierre et Marie Curie, Paris Cedex, France N Forteath Inspection Head Wharf, TAS, Australia J D Gage Scottish Association for Marine Science, Oban, UK S M Garcia Food and Agriculture Organization of the United Nations, Rome, Italy
G Ito University of Hawaii at Manoa, Honolulu, HI, USA J Jacoby Woods Hole Oceanographic Institution, Woods Hole, MA, USA M J Kaiser Bangor University, Bangor, UK A E S Kemp University of Southampton, Southampton Oceanography Centre, Southampton, UK W M Kemp University of Maryland Center for Environmental Science, Cambridge, MD, USA V S Kennedy University of Maryland, Cambridge, MD, USA P F Kingston Heriot-Watt University, Edinburgh, UK G L Kooyman University of California San Diego, CA, USA
C Garrett University of Victoria, VIC, Canada
W Krijgsman University of Utrecht, Utrecht, The Netherlands
R N Gibson Scottish Association for Marine Science, Argyll, Scotland
J B Kristoffersen University of Bergen, Bergen, Norway
M Gochfeld Environmental and Community Medicine, Piscataway, NJ, USA
K Lambeck Australian National University, Canberra, ACT, Australia
H O Halvorson University of Massachusetts Boston, Boston, MA, USA
R S Lampitt University of Southampton, Southampton, UK
B U Haq Vendome Court, Bethesda, MD, USA
M Landry University of Hawaii at Manoa, Department of Oceanography, Honolulu, HI, USA
G R Harbison Woods Hole Oceanographic Institution, Woods Hole, MA, USA
C G Langereis University of Utrecht, Utrecht, The Netherlands A Lascaratos University of Athens, Athens, Greece
R M Haymon University of California, CA, USA
S Leibovich Cornell University, Ithaca, NY, USA
D L Hebert University of Rhode Island, RI, USA J E Heyning The Natural History Museum of Los Angeles County, Los Angeles, CA, USA P Hoagland Woods Hole Oceanographic Institution, Woods Hole, MA, USA
W G Leslie Harvard University, Cambridge, MA, USA C Llewellyn Plymouth Marine Laboratory, Plymouth, UK R A Lutz Rutgers University, New Brunswick, NJ, USA
(c) 2011 Elsevier Inc. All Rights Reserved.
xx
Contributors
K C Macdonald Department of Geological Sciences and Marine Sciences Institute, University of California, Santa Barbara, CA, USA F T Mackenzie University of Hawaii, Honolulu, HI, USA L P Madin Woods Hole Oceanographic Institution, Woods Hole, MA, USA M Maslin University College London, London, UK G A Maul Florida Institute of Technology, Melbourne, FL, USA M McNutt MBARI, Moss Landing, CA, USA M G McPhee McPhee Research Company, Naches, WA, USA A D Mclntyre University of Aberdeen, Aberdeen, UK J Mienert University of Tromsø, Tromsø, Norway G E Millward University of Plymouth, Plymouth, UK H Momma Japan Marine Science & Technology Center, Japan J H Morison University of Washington, Seattle, WA, USA A E Mulligan Woods Hole Oceanographic Institution, Woods Hole, MA, USA
J E Petersen Oberlin College, Oberlin, OH, USA M Phillips Network of Aquaculture Centres in Asia-Pacific (NACA), Bangkok, Thailand B Qiu University of Hawaii at Manoa, Hawaii, USA F Quezada Biotechnology Center of Excellence Corporation, Waltham, MA, USA N N Rabalais Louisiana Universities Marine Consortium, Chauvin, LA, USA R D Ray NASA Goddard Space Flight Center, Greenbelt, MD, USA M R Reeve National Science Foundation, Arlington VA, USA R R Reeves Okapi Wildlife Associates, QC, Canada A Reyes-Prieto University of Iowa, Iowa City, IA, USA P B Rhines University of Washington,Seattle, WA, USA A R Robinson Harvard University, Cambridge, MA, USA H T Rossby University of Rhode Island, Kingston, RI, USA H M Rozwadowski Georgia Institute of Technology, Atlanta, Georgia, USA
W Munk University of California San Diego, La Jolla, CA, USA
A G V Salvanes University of Bergen, Bergen, Norway
E J Murphy British Antarctic Survey, Marine Life Sciences Division, Cambridge, UK
R Schlitzer Alfred Wegener Institute for Polar and Marine Research, Bremerhaven, Germany
P D Naidu National Institute of Oceanography, Dona Paula, India
M E Schumacher Woods Hole Oceanographic Institution, Woods Hole, MA, USA
N Niitsuma Shizuoka University, Shizuoka, Japan
M I Scranton State University of New York, Stony Brook, NY, USA
D B Olson University of Miami, Miami, FL, USA G-A Paffenho¨fer Skidaway Institute of Oceanography, Savannah, GA, USA C Paris University of Miami, Miami, FL, USA M R Perfit Department of Geological Sciences, University of Florida, Gainsville, FL, USA
K Sherman Narragansett Laboratory, Narragansett, RI, USA M D Spalding UNEP World Conservation Monitoring Centre and Cambridge Coastal Research Unit, Cambridge, UK J Sprintall University of California San Diego, La Jolla, CA, USA J H Steele Woods Hole Oceanographic Institution, MA, USA
(c) 2011 Elsevier Inc. All Rights Reserved.
Contributors C A Stein University of Illinois at Chicago, Chicago, IL, USA
S M Van Parijs Norwegian Polar Institute, Tromsø, Norway
C Stickley University College London, London, UK
L M Ver University of Hawaii, Honolulu, HI, USA
U R Sumaila University of British Columbia, Vancouver, BC, Canada
F J Vine University of East Anglia, Norwich, UK
S Takagawa Japan Marine Science & Technology Center, Japan
K L Von Damm University of New Hampshire, Durham, NH, USA
P K Taylor Southampton Oceanography Centre, Southampton, UK
R P Von Herzen Woods Hole Oceanographic Institution, Woods Hole, MA, USA
A Theocharis National Centre for Marine Research (NCMR), Hellinikon, Athens, Greece
xxi
D Wartzok Florida International University, Miami, FL, USA
P C Ticco Massachusetts Maritime Academy, Buzzards Bay, MA, USA R P Trask Woods Hole Oceanographic Institution, Woods Hole, MA, USA
W F Weeks Portland, OR, USA R A Weller Woods Hole Oceanographic Institution, Woods Hole, MA, USA
A W Trites University of British Columbia, British Columbia, Canada
J A Whitehead Woods Hole Oceanographic Institution, Woods Hole, MA, USA
A Turner University of Plymouth, Plymouth, UK
J C Wiltshire University of Hawaii, Manoa, Honolulu, HA, USA
P L Tyack Woods Hole Oceanographic Institution, Woods Hole, MA, USA
C Woodroffe University of Wollongong, Wollongong, NSW, Australia
G J C Underwood University of Essex, Colchester, UK
C Wunsch Massachusetts Institute of Technology, Cambridge, MA, USA
C L Van Dover The College of William and Mary, Williamsburg, VA, USA
H S Yoon University of Iowa, Iowa City, IA, USA
Volume 4 A Alldredge University of California, Santa Barbara, CA, USA D M Anderson Woods Hole Oceanographic Institution, Woods Hole, MA, USA O R Anderson Columbia University, Palisades, NY, USA
J M Bewers Bedford Institute of Oceanography, Dartmouth, NS, Canada N V Blough University of Maryland, College Park, MD, USA W Bonne
P G Baines CSIRO Atmospheric Research, Aspendale, VIC, Australia
Federal Public Service Health, Food Chain Safety and Environment, Brussels, Belgium
J M Baker Clock Cottage, Shrewsbury, UK
University of Cape Town, Cape Town, Republic of South Africa
J G Bellingham Monterey Bay Aquarium Research Institute, Moss Landing, CA, USA
R D Brodeur
G M Branch
Northwest Fisheries Science Center, Newport, OR, USA
(c) 2011 Elsevier Inc. All Rights Reserved.
xxii
Contributors
H Burchard Baltic Sea Research Institute Warnemu¨nde, Warnemu¨nde, Germany P H Burkill Plymouth Marine Laboratory, West Hoe, Plymouth, UK Francois Carlotti C.N.R.S./Universite´ Bordeaux 1, Arachon, France K L Casciotti Woods Hole Oceanographic Institution, Woods Hole, MA, USA
J S Grayw University of Oslo, Oslo, Norway A G Grottoli University of Pennsylvania, Philadelphia, PA, USA N Gruber Institute of Biogeochemistry and Pollutant Dynamics, ETH Zurich, Switzerland K C Hamer University of Durham, Durham, UK D Hammond University of Southern California, Los Angeles, CA, USA
A Clarke British Antarctic Survey, Cambridge, UK
W W Hay Christian-Albrechts University, Kiel, Germany
M B Collins National Oceanography Centre, Southampton, UK
J W Heath Coastal Fisheries Institute, CCEER Louisiana State University, Baton Rouge, LA, USA
J J Cullen Department of Oceanography, Halifax, NS, Canada D H Cushing Lowestoft, Suffolk, UK
D Hedgecock University of Southern California, Los Angeles, CA, USA C Hemleben Tu¨bingen University, Tu¨bingen, Germany
K L Denman University of Victoria, Victoria, BC, Canada S C Doney Woods Hole Oceanographic Institution, Woods Hole, MA, USA
T D Herbert Brown University, Providence, RI, USA I Hewson University of California Santa Cruz, Santa Cruz, CA, USA
J F Dower University of British Columbia, Vancouver, BC, Canada
Richard Hey University of Hawaii at Manoa, Honolulu, HI, USA
K Dysthe University of Bergen, Bergen, Norway
P Hoagland Woods Hole Oceanographic Institution, Woods Hole, MA, USA
H N Edmonds University of Texas at Austin, Port Aransas, TX, USA
N Hoepffner Institute for Environment and Sustainability, Ispra, Italy
L Føyn Institute of Marine Research, Bergen, Norway
M Hood Intergovernmental Oceanographic Commission, Paris, France
J Fuhrman University of Southern California, Los Angeles, CA, USA
M J Howarth Proudman Oceanographic Laboratory, Wirral, UK
C P Gallienne Plymouth Marine Laboratory, West Hoe, Plymouth, UK
M Huber Purdue University, West Lafayette, IN, USA
E Garel CIACOMAR, Algarve University, Faro, Portugal
J W Hurrell National Center for Atmospheric Research, Boulder, CO, USA
D M Glover Woods Hole Oceanographic Institution, Woods Hole, MA, USA S L Goodbred Jr State University of New York, Stony Brook, NY, USA J D M Gordon Scottish Association for Marine Science, Oban, Argyll, UK
D R Jackett CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia R A Jahnke Skidaway Institute of Oceanography, Savannah, GA, USA w
Deceased.
(c) 2011 Elsevier Inc. All Rights Reserved.
Contributors
xxiii
A Jarre University of Cape Town, Cape Town, South Africa
A F Michaels University of Southern California, Los Angeles, CA, USA
J Joseph La Jolla, CA, USA
J D Milliman College of William and Mary, Gloucester, VA, USA
D M Karl University of Hawaii at Manoa, Honolulu, HI, USA
C D Mobley Sequoia Scientific, Inc., WA, USA
K L Karsh Princeton University, Princeton, NJ, USA
M M Mullinw Scripps Institution of Oceanography, La Jolla, CA, USA
J Karstensen Universita¨t Kiel (IFM-GEOMAR), Kiel, Germany
P Mu¨ller University of Hawaii, Honolulu, HI, USA
R M Key Princeton University, Princeton, NJ, USA
L A Murray The Centre for Environment, Fisheries and Aquaculture Sciences, Lowestoft, UK
P D Killworth Southampton Oceanography Centre, Southampton, UK B Klinger Center for Ocean-Land-Atmosphere Studies (COLA), Calverton, MD, USA H E Krogstad NTNU, Trondheim, Norway I Laing Centre for Environment Fisheries and Aquaculture Science, Weymouth, UK
T Nagai Tokyo University of Marine Science and Technology, Tokyo, Japan K H Nisancioglu Bjerknes Centre for Climate Research, University of Bergen, Bergen, Norway Y Nozakiw University of Tokyo, Tokyo, Japan
G F Lane-Serff University of Manchester, Manchester, UK
K J Orians The University of British Columbia, Vancouver, BC, Canada
A Longhurst Place de I’Eglise, Cajarc, France
C A Paulson Oregon State University, Corvallis, OR, USA
R Lukas University of Hawaii at Manoa, Hawaii, USA
W G Pearcy Oregon State University, Corvallis, OR, USA
M Lynch University of California Santa Barbara, Santa Barbara, CA, USA
W S Pegau Oregon State University, Corvallis, OR, USA
M Macleod World Wildlife Fund, Washington, DC, USA E Maran˜o´n University of Vigo, Vigo, Spain S Martin University of Washington, Seattle, WA, USA S M Masutani University of Hawaii at Manoa, Honolulu, HI, USA I N McCave University of Cambridge, Cambridge, UK T J McDougall CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia C L Merrin The University of British Columbia, Vancouver, BC, Canada
T Platt Dalhousie University, NS, Canada J J Polovina National Marine Fisheries Service, Honolulu, HI, USA D Quadfasel Niels Bohr Institute, Copenhagen, Denmark J A Raven Biological Sciences, University of Dundee, Dundee, UK G E Ravizza Woods Hole Oceanographic Institution, Woods Hole, MA, USA A J Richardson University of Queensland, St. Lucia, QLD, Australia M Rubega University of Connecticut, Storrs, CT, USA w
Deceased.
(c) 2011 Elsevier Inc. All Rights Reserved.
xxiv
Contributors
K C Ruttenberg Woods Hole Oceanographic Institution, Woods Hole, MA, USA
K K Turekian Yale University, New Haven, CT, USA T Tyrrell University of Southampton, Southampton, UK
A G V Salvanes University of Bergen, Bergen, Norway
O Ulloa Universidad de Concepcio´n, Concepcio´n, Chile
S Sathyendranath Dalhousie University, NS, Canada
C M G Vivian The Centre for Environment, Fisheries and Aquaculture Sciences, Lowestoft, UK
R Schiebel Tu¨bingen University, Tu¨bingen, Germany F B Schwing NOAA Fisheries Service, Pacific Grove, CA, USA
J J Walsh University of South Florida, St. Petersburg, FL, USA
M P Seki National Marine Fisheries Service, Honolulu, HI, USA
R M Warwick Plymouth Marine Laboratory, Plymouth, UK
L J Shannon Marine and Coastal Management, Cape Town, South Africa
N C Wells Southampton Oceanography Centre, Southampton, UK
K Shepherd Institute of Ocean Sciences, Sidney, BC, Canada
J A Whitehead Woods Hole Oceanographic Institution, Woods Hole, MA, USA
D Siegel-Causey Harvard University, Cambridge MA, USA D M Sigman Princeton University, Princeton, NJ, USA A Soloviev Nova Southeastern University, FL, USA J H Steele Woods Hole Oceanographic Institution, MA, USA P K Takahashi University of Hawaii at Manoa, Honolulu, HI, USA L D Talley Scripps Institution of Oceanography, La Jolla, CA, USA E Thomas Yale University, New Haven, CT, USA J R Toggweiler NOAA, Princeton, NJ, USA
M Wilkinson Heriot-Watt University, Edinburgh, UK R G Williams University of Liverpool, Oceanography Laboratories, Liverpool, UK C A Wilson III Department of Oceanography and Coastal Sciences, and Coastal Fisheries Institute, CCEER Louisiana State University, Baton Rouge, LA, USA H Yamazaki Tokyo University of Marine Science and Technology, Tokyo, Japan B deYoung Memorial University, St. John’s, NL, Canada G Zibordi Institute for Environment and Sustainability, Ispra, Italy
Volume 5 D G Ainley H.T. Harvey Associates, San Jose CA, USA W Alpers University of Hamburg, Hamburg, Germany J R Apelw Global Ocean Associates, Silver Spring, MD, USA w
Deceased.
A B Baggeroer Massachusetts Institute of Technology, Cambridge, MA, USA L T Balance NOAA-NMFS, La Jolla, CA, USA R Batiza Ocean Sciences, National Science Foundation, VA, USA W H Berger Scripps Institution of Oceanography, La Jolla, CA, USA
(c) 2011 Elsevier Inc. All Rights Reserved.
Contributors J L Bodkin US Geological Survey, AK, USA
I Everson British Antarctic Survey Cambridge, UK
I L Boyd Natural Environment Research Council, Cambridge, UK
I Fer University of Bergen, Bergen, Norway
A C Brown University of Cape Town, Cape Town, Republic of South Africa
M Fieux Universite´-Pierre et Marie Curie, Paris, France
xxv
J Burger Rutgers University, Piscataway, NJ, USA
R A Flather Proudman Oceanographic Laboratory, Bidston Hill, Prenton, UK
C J Camphuysen Netherlands Institute for Sea Research, Texel, The Netherlands
G S Giese Woods Hole Oceanographic Institution, Woods Hole, MA, USA
D C Chapman Woods Hole Oceanographic Institution, Woods Hole, MA, USA
J M Gregory Hadley Centre, Berkshire, UK
R E Cheney Laboratory for Satellite Altimetry, Silver Spring, Maryland, USA T Chopin University of New Brunswick, Saint John, NB, Canada J A Church Antarctic CRC and CSIRO Marine Research, TAS, Australia J K Cochran State University of New York, Stony Brook, NY, USA P Collar Southampton Oceanography Centre, Southampton, UK R J Cuthbert University of Otago, Dunedin, New Zealand L S Davis University of Otago, Dunedin, New Zealand K L Denman University of Victoria, Victoria BC, Canada R P Dinsmore Woods Hole Oceanographic Institution, Woods Hole, MA, USA G J Divoky University of Alaska, Fairbanks, AK, USA
S M Griffies NOAA/GFDL, Princeton, NJ, USA G Griffiths Southampton Oceanography Centre, Southampton, UK A Harding University of California, San Diego, CA, USA W S Holbrook University of Wyoming, Laramie, WY, USA G L Hunt, Jr University of Washington, Seattle, WA, USA and University of California, Irvine, CA, USA P Hutchinson North Atlantic Salmon Conservation Organization, Edinburgh, UK K B Katsaros Atlantic Oceanographic and Meteorological Laboratory, NOAA, Miami, FL, USA H L Kite-Powell Woods Hole Oceanographic Institution, Woods Hole, MA, USA M A Kominz Western Michigan University, Kalamazoo, MI, USA
L M Dorman University of California, San Diego, La Jolla, CA, USA
R G Kope Northwest Fisheries Science Center, Seattle, WA, USA
J F Dower University of British Columbia, Vancouver, BC, Canada
G S E Lagerloef Earth and Space Research, Seattle, WA, USA
J B Edson Woods Hole Oceanographic Institution, Woods Hole, MA, USA
L M Lairdw Aberdeen University, Aberdeen, UK
T I Eglinton Woods Hole Oceanographic Institution, Woods Hole, MA, USA
M Leppa¨ranta University of Helsinki, Helsinki, Finland w
Deceased.
(c) 2011 Elsevier Inc. All Rights Reserved.
xxvi
Contributors
E J Lindstrom NASA Science Mission Directorate, Washington, DC, USA
C T Roman University of Rhode Island, Narragansett, RI, USA
A K Liu NASA Goddard Space Flight Center, Greenbelt, MD, USA
M Sawhney University of New Brunswick, Saint John, NB, Canada
C R McClain NASA Goddard Space Flight Center, Greenbelt, MD, USA
G Shanmugam The University of Texas at Arlington, Arlington, TX, USA
D J McGillicuddy Jr Woods Hole Oceanographic Institution, Woods Hole, MA, USA
J Sharples Proudman Oceanographic Laboratory, Liverpool, UK
W K Melville Scripps Institution of Oceanography, La Jolla CA, USA
J H Simpson Bangor University, Bangor, UK R K Smedbol Dalhousie University, Halifax, NS, Canada
D Mills Atlantic Salmon Trust, UK
L B Spear H.T. Harvey Associates, San Jose, CA, USA
P J Minnett University of Miami, Miami, FL, USA W A Montevecchi Memorial University of Newfoundland, NL, Canada W S Moore University of South Carolina, Columbia, SC, USA S J Morreale Cornell University, Ithaca, NY, USA K W Nicholls British Antarctic Survey, Cambridge, UK T J O’Shea Midcontinent Ecological Science Center, Fort Collins, CO, USA T E Osterkamp University of Alaska, Alaska, AK, USA F V Paladino Indiana-Purdue University at Fort Wayne, Fort Wayne, IN, USA C L Parkinson NASA Goddard Space Flight Center, Greenbelt, MD, USA A Pearson Woods Hole Oceanographic Institution, Woods Hole, MA, USA J T Potemra SOEST/IPRC, University of Hawaii, Honolulu, HI, USA J A Powell Florida Marine Research Institute, St Petersburg, FL, USA T Qu SOEST/IPRC, University of Hawaii, Honolulu, HI, USA
R L Stephenson St. Andrews Biological Station, St. Andrews, NB, Canada J M Teal Woods Hole Oceanographic Institution, Rochester, MA, USA K K Turekian Yale University, New Haven, CT, USA P Wadhams University of Cambridge, Cambridge, UK W F Weeks Portland, OR, USA G Wefer Universita¨t Bremen, Bremen, Germany W S Wilson NOAA/NESDIS, Silver Spring, MD, USA M Windsor, North Atlantic Salmon Conservation Organization, Edinburgh, UK S Y Wu NASA Goddard Space Flight Center, Greenbelt, MD, USA L Yu Woods Hole Oceanographic Institution, Woods Hole, MA, USA
(c) 2011 Elsevier Inc. All Rights Reserved.
Contributors
xxvii
Volume 6 A V Babanin Swinburne University of Technology, Melbourne, VIC, Australia
S E Humphris Woods Hole Oceanographic Institution, Woods Hole, MA, USA
R T Barber Duke University Marine Laboratory, Beaufort, NC, USA
W J Jenkins University of Southampton, Southampton, UK
J Bartram World Health Organization, Geneva, Switzerland
D R B Kraemer The Johns Hopkins University, Baltimore, MD, USA
A Beckmann Alfred-Wegener-Institut fu¨r Polar- und Meeresforschung, Bremerhaven, Germany M C Benfield Louisiana State University, Baton Rouge, LA, USA P S Bogden Maine State Planning Office, Augusta, ME, USA J A T Bye The University of Melbourne, Melbourne, VIC, Australia M F Cronin NOAA Pacific Marine Environmental Laboratory, Seattle, WA, USA A R J David Bere Alston, Devon, UK W Deuser Woods Hole Oceanographic Institution, Woods Hole, MA, USA J Donat Old Dominion University, Norfolk, VA, USA C Dryden Old Dominion University, Norfolk, VA, USA A Dufour United States Environmental Protection Agency, OH, USA C A Edwards University of Connecticut, Groton, CT, USA W J Emery University of Colorado, Boulder, CO, USA E Fahrbach Alfred-Wegener-Institut fu¨r Polar- und Meeresforschung, Bremerhaven, Germany
S Krishnaswami Physical Research Laboratory, Ahmedabad, India E L Kunze University of Washington, Seattle, WA, USA T E L Langford University of Southampton, Southampton, UK J R Ledwell Woods Hole Oceanographic Institution, Woods Hole, MA, USA P L-F Liu Cornell University, Ithaca, NY, USA M M R van der Loeff Alfred-Wegener-Institut fu¨r Polar und Meereforschung Bremerhaven, Germany R Lueck University of Victoria, Victoria, BC, Canada J E Lupton Hatfield Marine Science Center, Newport, OR, USA L P Madin Woods Hole Oceanographic Institution, Woods Hole, MA, USA M E McCormick The Johns Hopkins University, Baltimore, MD, USA M G McPhee McPhee Research Company, Naches, WA, USA J H Middleton The University of New South Wales, Sydney, NSW, Australia P J Minnett University of Miami, Miami, FL, USA E C Monahan University of Connecticut at Avery Point, Groton, CT, USA
A M Gorlov Northeastern University, Boston, Massachusetts, USA
C Moore WET Labs Inc., Philomath, OR, USA
I Helmond CSIRO Marine Research, TAS, Australia
J H Morison University of Washington, Seattle, WA, USA
R A Holman Oregon State University, Corvallis, OR, USA
J N Moum Oregon State University, Corvallis, OR, USA
(c) 2011 Elsevier Inc. All Rights Reserved.
xxviii
Contributors
N S Oakey Bedford Institute of Oceanography, Dartmouth, NS, Canada D T Pugh University of Southampton, Southampton, UK D L Rudnick University of California, San Diego, CA, USA H Salas CEPIS/HEP/Pan American Health Organization, Lima, Peru L K Shay University of Miami, Miami, FL, USA W D Smyth Oregon State University, Corvallis, OR, USA J Sprintall University of California San Diego, La Jolla, CA, USA
L St. Laurrent University of Victoria, Victoria, BC, Canada W G Sunda National Ocean Service, NOAA, Beaufort, NC, USA M Tomczak Flinders University of South Australia, Adelaide, SA, Australia A J Watson University of East Anglia, Norwich, UK P H Wiebe Woods Hole Oceanographic Institution, Woods Hole, MA, USA P F Worcester University of California at San Diego, La Jolla, CA, USA
(c) 2011 Elsevier Inc. All Rights Reserved.
Contents Volume 1 Abrupt Climate Change
S Rahmstorf
1
Absorbance Spectroscopy for Chemical Sensors Abyssal Currents
R Narayanaswamy, F Sevilla, III
W Zenk
Accretionary Prisms
15
J C Moore
31
Acoustic Measurement of Near-Bed Sediment Transport Processes Acoustic Noise
Acoustic Scintillation Thermography Acoustics In Marine Sediments
K G Foote
62
P A Rona, C D Jones
71
T Akal
75
P N Mikhalevsky
Acoustics, Deep Ocean
92
W A Kuperman
101
F B Jensen
112
Acoustics, Shallow Water R Chester
Agulhas Current
120
J R E Lutjeharms
Aircraft Remote Sensing
128
L W Harding Jr, W D Miller, R N Swift, C W Wright
Air–Sea Gas Exchange
38 52
Acoustic Scattering by Marine Organisms
Aeolian Inputs
P D Thorne, P S Bell
I Dyer
Acoustics, Arctic
7
B Ja¨hne
138 147
Air–Sea Transfer: Dimethyl Sulfide, COS, CS2, NH4, Non-Methane Hydrocarbons, Organo-Halogens J W Dacey, H J Zemmelink
157
Air–Sea Transfer: N2O, NO, CH4, CO
163
Alcidae
C S Law
T Gaston
171
Antarctic Circumpolar Current Antarctic Fishes
S R Rintoul
I Everson
191
Anthropogenic Trace Elements in the Ocean Antifouling Materials
E A Boyle
211
W Seaman, W J Lindberg
234
R A Duce
Atmospheric Transport and Deposition of Particulate Material to the Oceans R Arimoto Authigenic Deposits
226
S G Philander
Atmospheric Input of Pollutants
Baleen Whales
203
B Rudels
Atlantic Ocean Equatorial Currents
Bacterioplankton
195
D J Howell, S M Evans
Arctic Ocean Circulation Artificial Reefs
178
G M McMurtry H W Ducklow
J L Bannister
238 J M Prospero, 248 258 269 276
(c) 2011 Elsevier Inc. All Rights Reserved.
xxix
xxx
Contents
Baltic Sea Circulation Bathymetry
W Krauss
288
D Monahan
297
Beaches, Physical Processes Affecting Benguela Current
Benthic Foraminifera
316 D J Wildish
328
A J Gooday
Benthic Organisms Overview
336
P F Kingston
348
P L Tyack
357
Biogeochemical Data Assimilation
E E Hofmann, M A M Friedrichs
Biological Pump and Particle Fluxes Bioluminescence
Bioturbation
S Honjo
376
A Morel
385
D H Shull
Black Sea Circulation
395
G I Shapiro
Bottom Water Formation
401
A L Gordon
415
Brazil and Falklands (Malvinas) Currents
A R Piola, R P Matano
Breaking Waves and Near-Surface Turbulence
J Gemmrich
D K Woolf
Calcium Carbonates
L C Peterson
E D Barton
Carbon Dioxide (CO2) Cycle
467
T Takahashi
Cenozoic Climate – Oxygen Isotope Evidence Cenozoic Oceans – Carbon Cycle Models
J D Wright L Franc¸ois, Y Godde´ris
R A Fine W R Martin
J W Farrington
Coastal Zone Management
514
539 551 563
F E Werner, B O Blanton
Coastal Topography, Human Impact on Coastal Trapped Waves
502
531
H Chamley
Coastal Circulation Models
495
524
Chemical Processes in Estuarine Sediments
Coccolithophores
E E Adams, K Caldeira
P Boyle
Chlorinated Hydrocarbons
477 487
Carbon Sequestration via Direct Injection into the Ocean
Clay Mineralogy
455
C A Carlson, N R Bates, D A Hansell, D K Steinberg
CFCs in the Ocean
431
445
B M Hickey, T C Royer
Canary and Portugal Currents
Cephalopods
422
439
California and Alaska Currents
Carbon Cycle
364 371
P J Herring, E A Widder
Bio-Optical Models
Bubbles
305
L V Shannon
Benthic Boundary Layer Effects
Bioacoustics
A D Short
D M Bush, O H Pilkey, W J Neal
J M Huthnance D R Godschalk
T Tyrrell, J R Young
(c) 2011 Elsevier Inc. All Rights Reserved.
572 581 591 599 606
Contents
Cold-Water Coral Reefs Conservative Elements
J M Roberts
615
D W Dyrssen
626
Continuous Plankton Recorders Copepods
A John, P C Reid
R Harris
Coral Reefs
630 640
Coral Reef and Other Tropical Fisheries Coral Reef Fishes
xxxi
V Christensen, D Pauly
M A Hixon
655
J W McManus
660
Corals and Human Disturbance Cosmogenic Isotopes
N J Pilcher
671
D Lal
678
Coupled Sea Ice–Ocean Models Crustacean Fisheries
651
A Beckmann, G Birnbaum
688
J W Penn, N Caputi, R Melville-Smith
699
CTD (Conductivity, Temperature, Depth) Profiler Current Systems in the Atlantic Ocean Current Systems in the Indian Ocean
A J Williams, III
L Stramma M Fieux, G Reverdin
Current Systems in the Southern Ocean
A L Gordon
Current Systems in the Mediterranean Sea
P Malanotte-Rizzoli
708 718 728 735 744
Volume 2 Data Assimilation in Models Deep Convection
A R Robinson, P F J Lermusiaux
J R N Lazier
Deep Submergence, Science of
13 D J Fornari
22
K Moran
37
Deep-Sea Drilling Methodology Deep-Sea Drilling Results
1
J G Baldauf
45
Deep-Sea Fauna
P V R Snelgrove, J F Grassle
55
Deep-Sea Fishes
J D M Gordon
67
Deep-Sea Ridges, Microbiology
A-L Reysenbach
73
Deep-Sea Sediment Drifts
D A V Stow
80
Demersal Species Fisheries
K Brander
90
Determination of Past Sea Surface Temperatures Differential Diffusion
A E Gargett
Dispersion from Hydrothermal Vents Diversity of Marine Species Dolphins and Porpoises
R W Schmitt, J R Ledwell
K R Helfrich
P V R Snelgrove R S Wells
Double-Diffusive Convection
98 114
Dispersion and Diffusion in the Deep Ocean
Drifters and Floats
M Kucera
R W Schmitt
P L Richardson
(c) 2011 Elsevier Inc. All Rights Reserved.
122 130 139 149 162 171
xxxii
Contents
Dynamics of Exploited Marine Fish Populations East Australian Current
M J Fogarty
G Cresswell
179 187
Economics of Sea Level Rise
R S J Tol
197
Ecosystem Effects of Fishing
S J Hall
201
Eels
J D McCleave
208
Effects of Climate Change on Marine Mammals Ekman Transport and Pumping
T K Chereskin, J F Price
El Nin˜o Southern Oscillation (ENSO)
Electrical Properties of Sea Water
Energetics of Ocean Mixing
228
S G Philander
R D Prien
Elemental Distribution: Overview
Y Nozaki
255
A C Naveira Garabato
261
Eutrophication
271
J M Klymak, J D Nash
Estuarine Circulation
288
K Dyer
299
V N de Jonge, M Elliott
Evaporation and Humidity
Fiord Circulation
306
K Katsaros
Exotic Species, Introduction of Expendable Sensors
241 247
w
A V Fedorov, J N Brown
Estimates of Mixing
218 222
K E Trenberth
El Nin˜o Southern Oscillation (ENSO) Models
Equatorial Waves
I Boyd, N Hanson
324
D Minchin
332
J Scott
345
A Stigebrandt
353
Fiordic Ecosystems
K S Tande
359
Fish Ecophysiology
J Davenport
367
Fish Feeding and Foraging Fish Larvae
P J B Hart
E D Houde
Fish Locomotion
381
J J Videler
Fish Migration, Horizontal Fish Migration, Vertical
Fish Reproduction
Fish Vision
392
G P Arnold
402
J D Neilson, R I Perry
Fish Predation and Mortality
Fish Schooling
374
411
K M Bailey, J T Duffy-Anderson
J H S Blaxter
425
T J Pitcher
432
R H Douglas
445
Fish: Demersal Fish (Life Histories, Behavior, Adaptations) Fish: General Review
O A Bergstad
Q Bone
458 467
Fish: Hearing, Lateral Lines (Mechanisms, Role in Behavior, Adaptations to Life Underwater) A N Popper, D M Higgs w
417
Deceased.
(c) 2011 Elsevier Inc. All Rights Reserved.
476
Contents
Fisheries and Climate
K M Brander
Fisheries Economics
483
U R Sumaila, G R Munro
Fisheries Overview
491
M J Fogarty, J S Collie
Fisheries: Multispecies Dynamics Fishery Management
499
J S Collie
505
T P Smith, M P Sissenwine
Fishery Management, Human Dimension
513
D C Wilson, B J McCay
Fishery Manipulation through Stock Enhancement or Restoration Fishing Methods and Fishing Fleets Floc Layers
xxxiii
M D J Sayer
R Fonteyne
522 528 535
R S Lampitt
548
Florida Current, Gulf Stream, and Labrador Current Flow through Deep Ocean Passages Flows in Straits and Channels
P L Richardson
N G Hogg
554 564
D M Farmer
572
Fluid Dynamics, Introduction, and Laboratory Experiments
S A Thorpe
578
Fluorometry for Biological Sensing
D J Suggett, C M Moore
581
Fluorometry for Chemical Sensing
S Draxler, M E Lippitsch
589
Food Webs
A Belgrano, J A Dunne, J Bascompte
Forward Problem in Numerical Models Fossil Turbulence
596
M A Spall
604
C H Gibson
612
Volume 3 Gas Exchange in Estuaries
M I Scranton, M A de Angelis
Gelatinous Zooplankton
L P Madin, G R Harbison
General Circulation Models
Geomorphology
20
C G Langereis, W Krijgsman
C Woodroffe
Geophysical Heat Flow
C A Stein, R P Von Herzen
40 K Lambeck
C C Eriksen A D Mclntyre
Grabs for Shelf Benthic Sampling
67
P F Kingston
70
M McNutt
80
Groundwater Flow to the Coastal Ocean Habitat Modification
49 59
Global Marine Pollution
Gravity
25 33
Glacial Crustal Rebound, Sea Levels, and Shorelines Gliders
9
G R Ierley
Geomagnetic Polarity Timescale
1
A E Mulligan, M A Charette
M J Kaiser
Heat and Momentum Fluxes at the Sea Surface Heat Transport and Climate History of Ocean Sciences
88 99
P K Taylor
H L Bryden H M Rozwadowski
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105 114 121
xxxiv
Contents
Holocene Climate Variability Hydrothermal Vent Biota
M Maslin, C Stickley, V Ettwein R A Lutz
125 133
Hydrothermal Vent Deposits
R M Haymon
144
Hydrothermal Vent Ecology
C L Van Dover
151
Hydrothermal Vent Fauna, Physiology of
A J Arp
159
Hydrothermal Vent Fluids, Chemistry of
K L Von Damm
164
Hypoxia
N N Rabalais
172
Icebergs
D Diemand
181
Ice-Induced Gouging of the Seafloor Ice–Ocean Interaction
W F Weeks
J H Morison, M G McPhee
191 198
Ice Shelf Stability
C S M Doake
209
Igneous Provinces
M F Coffins, O Eldholm
218
Indian Ocean Equatorial Currents Indonesian Throughflow
M Fieux
J Sprintall
237
Inherent Optical Properties and Irradiance Internal Tidal Mixing Internal Tides
T D Dickey
W Munk
258
C Garrett
266
International Organizations Intertidal Fishes
M R Reeve
274
R N Gibson
Intra-Americas Sea
244 254
R D Ray
Internal Waves
Intrusions
226
280
G A Maul
286
D L Hebert
295
Inverse Modeling of Tracers and Nutrients
R Schlitzer
300
Inverse Models
C Wunsch
312
IR Radiometers
C J Donlon
319
K H Coale
331
Iron Fertilization Island Wakes Krill
E D Barton
343
E J Murphy
349
Kuroshio and Oyashio Currents
B Qiu
Laboratory Studies of Turbulent Mixing Lagoons
358 J A Whitehead
R S K Barnes
Lagrangian Biological Models Land–Sea Global Transfers
377 D B Olson, C Paris, R Cowen F T Mackenzie, L M Ver
Langmuir Circulation and Instability Large Marine Ecosystems
S Leibovich
K Sherman
Laridae, Sternidae, and Rynchopidae Law of the Sea
371
389 394 404 413
J Burger, M Gochfeld
P Hoagland, J Jacoby, M E Schumacher (c) 2011 Elsevier Inc. All Rights Reserved.
420 432
Contents
Leeuwin Current
G Cresswell, C M Domingues
Long-Term Tracer Changes Macrobenthos Magnetics
444
F von Blanckenburg
455
J D Gage
467
F J Vine
478
Manganese Nodules Mangroves
xxxv
D S Cronan
488
M D Spalding
496
Manned Submersibles, Deep Water
H Hotta, H Momma, S Takagawa
Manned Submersibles, Shallow Water
T Askew
505 513
Mariculture Diseases and Health
A E Ellis
519
Mariculture of Aquarium Fishes
N Forteath
524
Mariculture of Mediterranean Species Mariculture Overview
G Barnabe´, F Doumenge
M Phillips
537
Mariculture, Economic and Social Impacts Marine Algal Genomics and Evolution Marine Biotechnology
532
C R Engle
545
A Reyes-Prieto, H S Yoon, D Bhattacharya
H O Halvorson, F Quezada
552 560
Marine Chemical and Medicine Resources
S Ali, C Llewellyn
567
Marine Fishery Resources, Global State of
J Csirke, S M Garcia
576
Marine Mammal Diving Physiology
G L Kooyman
Marine Mammal Evolution and Taxonomy
J E Heyning
Marine Mammal Migrations and Movement Patterns Marine Mammal Overview
582
P J Corkeron, S M Van Parijs
P L Tyack
Marine Mammal Trophic Levels and Interactions Marine Mammals and Ocean Noise
A W Trites
Marine Policy Overview Marine Protected Areas Marine Silica Cycle
635 643
654 G-A Paffenho¨fer
656
P Hoagland, P C Ticco
664
P Hoagland, U R Sumaila, S Farrow D J DeMaster
R S Lampitt
Maritime Archaeology
622
651
J H Steele
Marine Plankton Communities
Mediterranean Sea Circulation
672 678 686
R D Ballard
Meddies and Sub-Surface Eddies
Meiobenthos
S K Hooker
A E S Kemp
Marine Mesocosms
615
628
R R Reeves
Marine Mammals: Sperm Whales and Beaked Whales
Marine Snow
P L Tyack
D Wartzok
Marine Mammals, History of Exploitation
596 605
Marine Mammal Social Organization and Communication
Marine Mats
589
H T Rossby A R Robinson, W G Leslie, A Theocharis, A Lascaratos
B C Coull, G T Chandler (c) 2011 Elsevier Inc. All Rights Reserved.
695 702 710 726
xxxvi
Contents
Mesocosms: Enclosed Experimental Ecosystems in Ocean Science Mesopelagic Fishes
J E Petersen, W M Kemp
A G V Salvanes, J B Kristoffersen
Mesoscale Eddies
748
P B Rhines
Metal Pollution
755
G E Millward, A Turner
Metalloids and Oxyanions
732
768
G A Cutter
776
Methane Hydrates and Climatic Effects
B U Haq
784
Methane Hydrate and Submarine Slides
J Mienert
790
Microbial Loops
M Landry
Microphytobenthos
799
G J C Underwood
807
Mid-Ocean Ridge Geochemistry and Petrology Mid-Ocean Ridge Seismic Structure Mid-Ocean Ridge Seismicity
M R Perfit
815
S M Carbotte
826
D R Bohnenstiehl, R P Dziak
Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology
837 K C Macdonald
Mid-Ocean Ridges: Mantle Convection and Formation of the Lithosphere Millennial-Scale Climate Variability
J T Andrews
Mineral Extraction, Authigenic Minerals Molluskan Fisheries Monsoons, History of Moorings
G Ito, R A Dunn
852 867 881
J C Wiltshire
890
V S Kennedy
899
N Niitsuma, P D Naidu
910
R P Trask, R A Weller
919
Volume 4 Nekton
W G Pearcy, R D Brodeur
Nepheloid Layers
1
I N McCave
Network Analysis of Food Webs
8 J H Steele
Neutral Surfaces and the Equation of State Nitrogen Cycle
19 T J McDougall, D R Jackett
D M Karl, A F Michaels
Nitrogen Isotopes in the Ocean Noble Gases and the Cryosphere Non-Rotating Gravity Currents North Atlantic Oscillation (NAO) North Sea Circulation
25 32
D M Sigman, K L Karsh, K L Casciotti
40
M Hood
55
P G Baines
59
J W Hurrell
65
M J Howarth
73
Nuclear Fuel Reprocessing and Related Discharges
H N Edmonds
82
Ocean Biogeochemistry and Ecology, Modeling of
N Gruber, S C Doney
89
Ocean Carbon System, Modeling of Ocean Circulation
S C Doney, D M Glover
N C Wells
105 115
Ocean Circulation: Meridional Overturning Circulation
J R Toggweiler
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126
Contents
Ocean Gyre Ecosystems
M P Seki, J J Polovina
Ocean Margin Sediments Ocean Ranching
132
S L Goodbred Jr
138
A G V Salvanes
146
R G Williams
156
Ocean Subduction
Ocean Thermal Energy Conversion (OTEC) Ocean Zoning
S M Masutani, P K Takahashi
M Macleod, M Lynch, P Hoagland
Offshore Sand and Gravel Mining Oil Pollution
Okhotsk Sea Circulation
E Garel, W Bonne, M B Collins
200
H Yamazaki, H Burchard, K Denman, T Nagai
Open Ocean Convection
A Soloviev, B Klinger
Open Ocean Fisheries for Deep-Water Species
Optical Particle Characterization
P H Burkill, C P Gallienne
265 272 274
R Lukas
287
E Thomas
295 W W Hay
Paleoceanography: Orbitally Tuned Timescales Paleoceanography: the Greenhouse World Particle Aggregation Dynamics Past Climate from Corals
T D Herbert
M Huber, E Thomas
A Alldredge
A G Grottoli
K L Denman, J F Dower
Pelagic Biogeography
A Longhurst
D H Cushing
Pelecaniformes
Peru–Chile Current System
C A Paulson, W S Pegau
J Karstensen, O Ulloa
319 330 338 348 356
379 385 393
K C Ruttenberg
Photochemical Processes
311
370
M Rubega
Phosphorus Cycle
303
364
D Siegel-Causey
Penetrating Shortwave Radiation
252 261
I Laing
Paleoceanography, Climate Models in
Phytobenthos
R A Jahnke
K K Turekian
Pacific Ocean Equatorial Currents
Phalaropes
243
G F Lane-Serff
Oysters – Shellfish Farming
Pelagic Fishes
234
K K Turekian
Oxygen Isotopes in the Ocean
Paleoceanography
226
J Joseph
Organic Carbon Cycling in Continental Margin Environments
Overflows and Cascades
208 218
J D M Gordon
Open Ocean Fisheries for Large Pelagic Species
Origin of the Oceans
182 191
L D Talley
One-Dimensional Models
167 174
J M Baker
Patch Dynamics
xxxvii
N V Blough
M Wilkinson
401 414 425
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xxxviii
Contents
Phytoplankton Blooms
D M Anderson
Phytoplankton Size Structure Plankton
M M Mullin
Plankton and Climate Plankton Viruses
432
E Maran˜o´n
445
w
453
A J Richardson
455
J Fuhrman, I Hewson
465
Platforms: Autonomous Underwater Vehicles Platforms: Benthic Flux Landers
J G Bellingham
R A Jahnke
485
Platinum Group Elements and their Isotopes in the Ocean
G E Ravizza
Plio-Pleistocene Glacial Cycles and Milankovitch Variability Polar Ecosystems
K H Nisancioglu
A Clarke
Pollution, Solids
494 504 514
C M G Vivian, L A Murray
Pollution: Approaches to Pollution Control Pollution: Effects on Marine Communities Polynyas
473
519
J S Grayw, J M Bewers R M Warwick
526 533
S Martin
540
Population Dynamics Models
Francois Carlotti
Population Genetics of Marine Organisms Pore Water Chemistry
546
D Hedgecock
556
D Hammond
Primary Production Distribution
563
S Sathyendranath, T Platt
572
Primary Production Methods
J J Cullen
578
Primary Production Processes
J A Raven
585
Procellariiformes
K C Hamer
590
Propagating Rifts and Microplates
Richard Hey
597
Protozoa, Planktonic Foraminifera
R Schiebel, C Hemleben
606
Protozoa, Radiolarians
O R Anderson
Radiative Transfer in the Ocean Radioactive Wastes Radiocarbon
C D Mobley
619
L Føyn
629
R M Key
637
Rare Earth Elements and their Isotopes in the Ocean Red Sea Circulation Redfield Ratio Refractory Metals
Y Nozaki
w
D Quadfasel
677
K J Orians, C L Merrin
Regime Shifts, Physical Forcing Regime Shifts: Methods of Analysis
653 666
T Tyrrell
Regime Shifts, Ecological Aspects
w
613
L J Shannon, A Jarre, F B Schwing F B Schwing B deYoung, A Jarre
Deceased.
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687 699 709 717
Contents
Regional and Shelf Sea Models
J J Walsh
Remote Sensing of Coastal Waters
Rigs and offshore Structures River Inputs
722
N Hoepffner, G Zibordi
Remotely Operated Vehicles (ROVs)
xxxix
732
K Shepherd
742
C A Wilson III, J W Heath
748
J D Milliman
754
Rocky Shores
G M Branch
762
Rogue Waves
K Dysthe, H E Krogstad, P Mu¨ller
770
Rossby Waves
P D Killworth
Rotating Gravity Currents
781
J A Whitehead
790
Volume 5 Salmon Fisheries, Atlantic
P Hutchinson, M Windsor
Salmon Fisheries, Pacific Salmonid Farming Salmonids
1
R G Kope
L M Laird
12
w
23
D Mills
29
Salt Marsh Vegetation
C T Roman
Salt Marshes and Mud Flats Sandy Beaches, Biology of Satellite Altimetry
39
J M Teal
43
A C Brown
49
R E Cheney
58
Satellite Oceanography, History, and Introductory Concepts J R Apel w
W S Wilson, E J Lindstrom, 65
Satellite Passive-Microwave Measurements of Sea Ice
C L Parkinson
80
Satellite Remote Sensing of Sea Surface Temperatures
P J Minnett
91
Satellite Remote Sensing SAR
A K Liu, S Y Wu
Satellite Remote Sensing: Ocean Color
C R McClain
Satellite Remote Sensing: Salinity Measurements Science of Ocean Climate Models Sea Ice
103
G S E Lagerloef
S M Griffies
P Wadhams
114 127 133 141
Sea Ice Dynamics
M Leppa¨ranta
159
Sea Ice: Overview
W F Weeks
170
Sea Level Change
J A Church, J M Gregory
179
Sea Level Variations Over Geologic Time Sea Otters
w
M A Kominz
J L Bodkin
185 194
Deceased.
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xl
Contents
Sea Surface Exchanges of Momentum, Heat, and Fresh Water Determined by Satellite Remote Sensing L Yu
202
Sea Turtles
212
F V Paladino, S J Morreale
Seabird Conservation
J Burger
Seabird Foraging Ecology Seabird Migration
220
L T Balance, D G Ainley, G L Hunt Jr
L B Spear
227 236
Seabird Population Dynamics
G L Hunt Jr
247
Seabird Reproductive Ecology
L S Davis, R J Cuthbert
251
Seabird Responses to Climate Change Seabirds and Fisheries Interactions
David G Ainley, G J Divoky C J Camphuysen
Seabirds as Indicators of Ocean Pollution Seabirds: An Overview Seals
265
W A Montevecchi
274
G L Hunt, Jr
279
I L Boyd
285
Seamounts and Off-Ridge Volcanism Seas of Southeast Asia
R Batiza
292
J T Potemra, T Qu
Seaweeds and their Mariculture Sediment Chronologies
305
T Chopin, M Sawhney
317
J K Cochran
327
Sedimentary Record, Reconstruction of Productivity from the Seiches
Seismic Structure
I Fer, W S Holbrook
L M Dorman
367
K B Katsaros
375
Sensors for Micrometeorological and Flux Measurements Shelf Sea and Shelf Slope Fronts
J B Edson
J Sharples, J H Simpson
H L Kite-Powell
Single Point Current Meters
T I Eglinton, A Pearson
P Collar, G Griffiths
436
Slides, Slumps, Debris Flows, and Turbidity Currents Small Pelagic Species Fisheries Small-Scale Patchiness, Models of
419 428
T J O’Shea, J A Powell G Shanmugam
R L Stephenson, R K Smedbol D J McGillicuddy Jr
Small-Scale Physical Processes and Plankton Biology
J F Dower, K L Denman
M Fieux
447 468 474 488 494
A B Baggeroer
Southern Ocean Fisheries
391
409
Single Compound Radiocarbon Measurements
Sonar Systems
382
401
R P Dinsmore
Somali Current
351 361
Sensors for Mean Meteorology
Shipping and Ports
333 344
A Harding
Seismology Sensors
Sirenians
G Wefer, W H Berger
D C Chapman, G S Giese
Seismic Reflection Methods for Study of the Water Column
Ships
257
504
I Everson
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Contents
Sphenisciformes
L S Davis
520
Stable Carbon Isotope Variations in the Ocean Storm Surges
K K Turekian
529
R A Flather
530
Sub Ice-Shelf Circulation and Processes Submarine Groundwater Discharge Sub-Sea Permafrost Surface Films
xli
K W Nicholls
541
W S Moore
551
T E Osterkamp
559
W Alpers
570
Surface Gravity and Capillary Waves
W K Melville
573
Volume 6 Temporal Variability of Particle Flux Thermal Discharges and Pollution
W Deuser
1
T E L Langford
10
Three-Dimensional (3D) Turbulence Tidal Energy Tides
W D Smyth, J N Moum
A M Gorlov
26
D T Pugh
Tomography
32
P F Worcester
Topographic Eddies Towed Vehicles
40
J H Middleton
57
I Helmond
Trace Element Nutrients
65
W G Sunda
Tracer Release Experiments
75
A J Watson, J R Ledwell
Tracers of Ocean Productivity
Transmissometry and Nephelometry Tritium–Helium Dating
87
W J Jenkins
93
Transition Metals and Heavy Metal Speciation
Tsunami
18
J Donat, C Dryden
100
C Moore
109
W J Jenkins
119
P L-F Liu
127
Turbulence in the Benthic Boundary Layer Turbulence Sensors
R Lueck, L St. Laurrent, J N Moum
N S Oakey
Under-Ice Boundary Layer
148
M G McPhee, J H Morison
Upper Ocean Heat and Freshwater Budgets Upper Ocean Mean Horizontal Structure Upper Ocean Mixing Processes
155
P J Minnett
163
M Tomczak
175
J N Moum, W D Smyth
185
Upper Ocean Structure: Responses to Strong Atmospheric Forcing Events Upper Ocean Time and Space Variability Upper Ocean Vertical Structure Upwelling Ecosystems
141
L K Shay
192
D L Rudnick
211
J Sprintall, M F Cronin
217
R T Barber
Uranium-Thorium Decay Series in the Oceans: Overview
225 M M R van der Loeff
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233
xlii
Contents
Uranium-Thorium Series Isotopes in Ocean Profiles Vehicles for Deep Sea Exploration
S E Humphris
Viral and Bacterial Contamination of Beaches Volcanic Helium
J Bartram, H Salas, A Dufour
285
Water Types and Water Masses
W J Emery
291
M E McCormick, D R B Kraemer
Waves on Beaches
267 277
E L Kunze
Wave Generation by Wind
244 255
J E Lupton
Vortical Modes
Wave Energy
S Krishnaswami
300
J A T Bye, A V Babanin
304
R A Holman
310
Weddell Sea Circulation
E Fahrbach, A Beckmann
318
Wet Chemical Analyzers
A R J David
326
Whitecaps and Foam
E C Monahan
Wind- and Buoyancy-Forced Upper Ocean Wind Driven Circulation
331 M F Cronin, J Sprintall
P S Bogden, C A Edwards
Zooplankton Sampling with Nets and Trawls
337 346
P H Wiebe, M C Benfield
355
Appendix 1. SI Units and Some Equivalences
373
Appendix 2. Useful Values
376
Appendix 3. Periodic Table of the Elements
377
Appendix 4. The Geologic Time Scale
378
Appendix 5. Properties of Seawater
379
Appendix 6. The Beaufort Wind Scale and Seastate
384
Appendix 7. Estimated Mean Oceanic Concentrations of the Elements
386
Appendix 8. Abbreviations
389
Appendix 9. Taxonomic Outline Of Marine Organisms
L P Madin
401
Appendix 10. Bathymetric Charts of the Oceans
412
Index
421
(c) 2011 Elsevier Inc. All Rights Reserved.
GAS EXCHANGE IN ESTUARIES M. I. Scranton, State University of New York, Stony Brook, NY, USA M. A. de Angelis, Humboldt State University, Arcata, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1113–1119, & 2001, Elsevier Ltd.
the gas in either the gas or liquid phase.) For gases that make up a large fraction of the atmosphere (O2, N2, Ar), pgas does not vary temporally or spatially. For trace atmospheric gases (carbon dioxide (CO2), methane (CH4), hydrogen (H2), nitrous oxide (N2O), and others), pgas may vary considerably geographically or seasonally, and may be affected by anthropogenic activity or local natural sources.
Introduction Many atmospherically important gases are present in estuarine waters in excess over levels that would be predicted from simple equilibrium between the atmosphere and surface waters. Since estuaries are defined as semi-enclosed coastal bodies of water that have free connections with the open sea and within which sea water is measurably diluted with fresh water derived from land drainage, they tend to be supplied with much larger amounts of organic matter and other compounds than other coastal areas. Thus, production of many gases is enhanced in estuaries relative to the rest of the ocean. The geometry of estuaries, which typically have relatively large surface areas compared to their depths, is such that flux of material (including gases) from the sediments, and fluxes of gases across the air/water interface, can have a much greater impact on the water composition than would be the case in the open ocean. Riverine and tidal currents are often quite marked, which also can greatly affect concentrations of biogenic gases.
Gas Solubility The direction and magnitude of the exchange of gases across an air/water interface are determined by the difference between the surface-water concentration of a given gas and its equilibrium concentration or gas solubility with respect to the atmosphere. The concentration of a specific gas in equilibrium with the atmosphere (Ceq) is given by Henry’s Law: Ceq ¼ pgas =KH
½1
where pgas is the partial pressure of the gas in the atmosphere, and KH is the Henry’s Law constant for the gas. Typically as temperature and salinity increase, gas solubility decreases. (Note that Henry’s Law and the Henry’s Law constant also may be commonly expressed in terms of the mole fraction of
Gas Exchange (Flux) Across the Air/ Water Interface The rate of gas exchange across the air/water interface for a specific gas is determined by the degree of disequilibrium between the actual surface concentration of a gas (Csurf) and its equilibrium concentration (Ceq), commonly expressed as R: R ¼ Csurf =Ceq
½2
If R ¼ 1, the dissolved gas is in equilibrium with the atmosphere and no net flux or exchange with the atmosphere occurs. For gases with Ro1, the dissolved gas is undersaturated with respect to the atmosphere and there is a net flux of the gas from the atmosphere to the water. For gases with R41, a net flux of the gas from the water to the atmosphere occurs.
Models of Gas Exchange The magnitude of the flux (F) in units of mass of gas per unit area per unit time across the air/water interface is a function of the magnitude of the difference between the dissolved gas concentration and its equilibrium concentration as given by Fick’s First Law of Diffusion: F ¼ kðCsurf Ceq Þ
½3
where k is a first order rate constant, which is a function of the specific gas and surface water conditions. The rate constant, k, also known as the transfer coefficient, has units of velocity and is frequently given as k ¼ D=z
½4
where D is the molecular diffusivity (in units of cm2 s1), and z is the thickness of the laminar layer at the
(c) 2011 Elsevier Inc. All Rights Reserved.
1
2
GAS EXCHANGE IN ESTUARIES
air/water interface, which limits the diffusion of gas across the interface. In aquatic systems, Csurf is easily measured by gas chromatographic analysis, and Ceq may be calculated readily if the temperature and salinity of the water are known. In order to determine the flux of gas (F) in or out of the water, the transfer coefficient (k) needs to be determined. The value of k is a function of the surface roughness of the water. In open bodies of water, wind speed is the main determinant of surface roughness. A number of studies have established a relationship between wind speed and either the transfer coefficient, k, or the liquid laminar layer thickness, z (Figure 1). The transfer coefficient is also related to the Schmidt number, Sc, defined as: Sc ¼ n=D
½5
where n is the kinematic viscosity of the water. In calmer waters, corresponding to wind speeds of o5 m s1, k is proportional to Sc2/3. At higher wind speeds, but where breaking waves are rare, k is proportional to S1/2. Therefore, if the transfer coefficient of one gas is known, the k value for any other gas can be determined as: Scn1 =Scn2
k1 =k2 ¼ Smooth surface regime
150
Rough surface regime
_ 2/ 3
Breaking wave (bubble) regime
_ 1/ 2
Kw Sc
Kw Sc
125
½6
where n ¼ the exponent. For short-term steady winds, the transfer coefficient for CO2 has been derived as kCO2 ¼ 0:31ðU10 Þ2 ðSc=600Þ0:5
½7
where U10 is the wind speed at a height of 10 m above the water surface. Eqns [6] and [7] can be used to estimate k for gases other than CO2. In most estuarine studies wind speeds are measured closer to the water surface. In such cases, the wind speed measured at 2 cm above the water surface can be approximated as 0.5U10. In restricted estuaries and tidally influenced rivers, wind speed may not be a good predictor of wind speed due to limited fetch or blockage of prevailing winds by shore vegetation. Instead, streambed-generated turbulence is likely to be more important than wind stress in determining water surface roughness. In such circumstances, the large eddy model may be used to approximate k as: k ¼ 1:46ðDul1 Þ1=2
½8
where u is the current velocity, and l is equivalent to the mean depth in shallow turbulent systems. Much of the reported uncertainty (and study to study variability in fluxes) is caused by differences in assumptions related to the transfer coefficient rather than to large changes in concentration of the gas in the estuary.
Less soluble gas (e.g O2)
Transfer velocity cm h
_1
Direct Gas Exchange Measurements 100
75
More soluble gas (e.g. CO2)
50
25
0.25 0
0
5
0.5 10
_1
0.75 U ms 15 U ms
_1
1 20
Figure 1 Idealized plot of transfer coefficient (Kw) as a function of wind speed (u) and friction velocity (u*). (Adapted with permission from Liss PS and Merlivat L (1985) Air–sea gas exchange rates. In: Buat-Menard P (ed.) The Role of Air–sea Exchange in Geochemical Cycling, p. 117. NATO ASI Series C, vol. 185. Dordrecht: Reidel.)
Gas exchange with the atmosphere for gases for which water column consumption and production processes are known can be estimated using a dissolved gas budget. Through time-series measurements of biological and chemical cycling, gas loss or gain across the air/water interface can be determined by difference. For example, in the case of dissolved O2, the total change in dissolved O2 concentration over time can be attributed to air/water exchange and biological processes. The contribution of biological processes to temporal changes in dissolved O2 may be estimated from concurrent measurements of phosphate and an assumed Redfield stoichiometry, and subtracted from the total change to yield an estimate of air/water O2 exchange. Gas fluxes also may be measured directly using a flux chamber that floats on the water surface. The headspace of the chamber is collected and analyzed over several time points to obtain an estimate of the net amount of gas crossing the air/water interface enclosed by the chamber. If the surface area enclosed
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GAS EXCHANGE IN ESTUARIES
by the chamber is known, a net gas flux can be determined. Some flux chambers are equipped with small fans that simulate ambient wind conditions. However, most chambers do not use fans and so do not take account of the effects of wind-induced turbulence on gas exchange. Despite this limitation, flux chambers are important tools for measuring gas exchange in environments (such as estuaries or streams) where limited fetch or wind breaks produced by shoreline vegetation make wind less important than current-induced turbulence in shallow systems. While flux chambers may alter the surface roughness and, hence, gas/exchange rates via diffusion, flux chambers or other enclosed gas capturing devices also are the best method for determining loss of gases across the air/water interface due to ebullition of gas bubbles from the sediment. Measurement of radon (Rn) deficiencies in the upper water column can be used to determine gas exchange coefficients and laminar layer thickness, which can then be applied to other gases using eqn [3]. In this method, gaseous 222Rn, produced by radioactive decay of 226Ra, is assumed to be in secular equilibrium within the water column. 222Rn is relatively short-lived and has an atmospheric concentration of essentially zero. Therefore, in nearsurface waters, a 222Rn deficiency is observed, due to flux of 222Rn across the air/water interface. The flux of 222Rn across the air/water interface can be determined by the depth-integrated difference in measured 222Rn and that which should occur based on the 226 Ra inventory. From this flux, the liquid laminar layer thickness, z, can be calculated. Other volatile tracers have been used in estuaries to determine gas exchange coefficients. These tracers, such as chlorofluorocarbons (CFCs) or sulfur hexafluoride (SF6), are synthetic compounds with no known natural source. Unlike 222Rn, these gases are stable in solution. These tracers may be added to the aquatic system and the decrease of the gas due to flux across the air/water interface monitored over time. In some estuaries, point sources of these compounds may exist and the decrease of the tracer with distance downstream may be used to determine k or z values for the estuary.
Individual Gases
plays an important role controlling atmospheric chemistry, including serving as a regulator of tropospheric ozone concentrations and a major sink for hydroxyl radicals in the stratosphere. Methane concentrations have been increasing annually at the rate of approximately 1–2% over the last two centuries. While the contribution of estuaries to the global atmospheric methane budget is small because of the relatively small estuarine global surface area, estuaries have been identified as sources of methane to the atmosphere and coastal ocean and contribute a significant fraction of the marine methane emissions to the atmosphere. Surface methane concentrations reported primarily from estuaries in North America and Europe range from 1 to 42000 nM throughout the tidal portion of the estuaries. Methane in estuarine surface waters is generally observed to be supersaturated (100% saturation B2–3 nM CH4) with R-values ranging from 0.7 to 1600 (Table 1). In general, estuarine methane concentrations are highest at the freshwater end of the estuary and decrease with salinity. This trend reflects riverine input as the major source of methane to most estuaries, with reported riverine methane concentrations ranging from 5 to 10 000 nM. Estuaries with large plumes have been observed to cause elevated methane concentrations in adjacent coastal oceans. In addition to riverine input, sources of methane to estuaries include intertidal flats and marshes, ground-water input, runoff
Table 1 Methane saturation values (R) and estimated fluxes to the atmosphere for US and European estuaries Geographical regiona
Rb
Flux CH4 (mmol m2 h1)
North Pacific coast, USA North Pacific coast, USA Columbia River, USA Hudson River, USA Tomales Bay, CA, USA Baltic Sea, Germany European Atlantic coast Atlantic coast, USA Pettaquamscutt Estuary, USA
3–290 1–550 78 18–376 2–37 10.5–1550 0.7–1580 n.a. 81–111
6.2–41.7c 3.6–8.3c 26.0c 4.7–40.4c 17.4–26.3c 9.4–15.6c 5.5c 102–1107c 0.8–14.2c 541–3375d
a
Methane (CH4)
Atmospheric methane plays an important role in the Earth’s radiative budget as a potent greenhouse gas, which is 3.7 times more effective than carbon dioxide in absorbing infrared radiation. Despite being present in trace quantities, atmospheric methane
3
For studies that report values for a single estuary, the major river feeding the estuary is provided. For studies that report values for more than one estuary, the oceanic area being fed by the estuaries is given. b R ¼ degree of saturation ¼ measured concentration/atmospheric equilibrium concentration. n.a. indicates values not available in reference. c Diffusive flux. d Ebullition (gas bubble) flux.
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4
GAS EXCHANGE IN ESTUARIES
from agricultural and pasture land, petroleum pollution, lateral input from exposed bank soils, wastewater discharge and emission from organic-rich anaerobic sediments, either diffusively or via ebullition (transport of gas from sediments as bubbles) and subsequent dissolution within the water column. Anthropogenically impacted estuaries or estuaries supplied from impacted rivers tend to be characterized by higher water-column methane concentrations relative to pristine estuarine systems. Seasonally, methane levels in estuaries are higher in summer compared with winter, primarily due to increased bacterial methane production (methanogenesis) in estuarine and riverine sediments. Methane can be removed from estuarine waters by microbial methane oxidation and emission to the atmosphere. Methane oxidation within estuaries can be quite rapid, with methane turnover times of o2 h to several days. Methane oxidation appears to be most rapid at salinities of less than about 6 (on the practical salinity units scale) and is strongly dependent on temperature, with highest oxidation rates occurring during the summer, when water temperatures are highest. Methane oxidation rates decrease rapidly with higher salinities. Methane diffusive fluxes to the atmosphere reported for estuaries (Table 1) fall within a narrow range of 3.6–41.7 mmol m2 h1 (2–16 mg CH4 m2 day1). Using a global surface area for estuaries of 1.4 106 km2 yields an annual emission of methane to the atmosphere from estuaries of 1–8 Tg y1. Because the higher flux estimates given in Table 1 generally were obtained close to the freshwater endmember of the estuary, the global methane estuarine emission is most likely within the range of 1– 3 Tg y1, corresponding to approximately 10% of the total global oceanic methane flux to the atmosphere, despite the much smaller global surface area of estuaries relative to the open ocean. Methane is also released to the atmosphere directly from anaerobic estuarine sediments via bubble formation and injection into the water column. Although small amounts of methane from bubbles may dissolve within the water column, the relatively shallow nature of the estuarine environment results in the majority of methane in bubbles reaching the atmosphere. The quantitative release of methane via this mechanism is difficult to evaluate due to the irregular and sporadic spatial and temporal extent of ebullition. Where ebullition occurs, the flux of methane to the atmosphere is considerably higher than diffusive flux (Table 1), but the areal extent of bubbling is relatively smaller than that of diffusive flux and, except in organic-rich stagnant areas such as tidal marshes, probably does not contribute
significantly to estuarine methane emissions to the atmosphere. Methane emission via ebullition has been observed to be at least partially controlled by tidal changes in hydrostatic pressure, with release of methane occurring at or near low tide when hydrostatic pressure is at a minimum. Nitrous Oxide (N2O)
Nitrous oxide is another important greenhouse gas that is present in elevated concentrations in estuarine environments. At present, N2O is responsible for about 5–6% of the anthropogenic greenhouse effect and is increasing in the atmosphere at a rate of about 0.25% per year. However, the role of estuaries in the global budget of the gas has only been addressed recently. Nitrous oxide is produced primarily as an intermediate during both nitrification (the oxidation of ammonium to nitrate) and denitrification (the reduction of nitrate, via nitrite and N2O, to nitrogen gas), although production by dissimilatory nitrate reduction to ammonium is also possible. In estuaries, nitrification and denitrification are both thought to be important sources. Factors such as the oxygen level in the estuary and the nitrate and ammonium concentrations of the water can influence which pathway is dominant, with denitrification dominating at very low, but non-zero, oxygen concentrations. Nitrous oxide concentrations are typically highest in the portions of the estuary closest to the rivers, and decrease with distance downstream. A number of workers have reported nitrous oxide maxima in estuarine waters at low salinities (o5–10 on the PSU scale), but this is not always the case. The turbidity maximum has been reported to be the site of maximum nitrification (presumably because of increased residence time for bacteria attached to suspended particulate matter, combined with elevated substrate (oxygen and ammonium)). Table 2 presents a summary of the data published for degree of saturation and air–estuary flux of nitrous oxide from a variety of estuaries, all of which are located in Europe and North America. Concentrations are commonly above that predicted from air–sea equilibrium, and estimates of fluxes range from 0.01 mmol m2 h1 to 5 mmol m2 h1. Ebullition is not important for nitrous oxide because it is much more soluble than methane. Researchers have estimated the size of the global estuarine source for N2O based on fluxes from individual estuaries multiplied by the global area occupied by estuaries to range from 0.22 Tg N2O y1 to 5.7 Tg y1 depending on the characteristics of the rivers studied. Independent estimates based on budgets of nitrogen
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GAS EXCHANGE IN ESTUARIES
Table 2 Nitrous oxide saturation values (R) and estimated fluxes to the atmosphere for US and European estuaries Estuary
R
Flux a N2O (mmol m2 h1)
Europe Gironde River Gironde River Oder River Elbe Scheldt Scheldt
1.1–1.6 E1.0–3.2 0.9–3.1 2.0–16 E1.0–31 E1.2–30
n.a. n.a. 0.014–0.165 n.a. 1.27–4.77 3.56
UK Colne Tamar Humber Tweed Mediterranean Amvraikos Gulf
0.9–13.6 1–3.3 2–40 0.96–1.1
0.9–1.1
1.3 0.41 1.8 E0
5
Table 3 Fluxes of carbon dioxide from estuaries in Europe and eastern USA Estuary
European rivers Northern Europe Scheldt estuary Portugal UK Clyde estuary East coast USA Hudson River (tidal freshwater) Georgia rivers
R
Fluxa CO2 (mmol m2 h1)
0.7–61.1 0.35–26.2 1.6–15.8 1.1–14.4 E0.7–1.8
1.0–31.7 4.2–50 10–31.7 4.4–10.4 n.a.
1.2–5.4
0.67–1.54
Slight supersaturation to 22.9
1.7–23
a n.a., insufficient data were available to permit calculation of this value.
0.043 7 0.0468
North-west USA Yaquina Bay Alsea River
1.0–4.0 0.9–2.4
0.165–0.699 0.047–0.72
East coast USA Chesapeake Bay Merrimack
0.9–1.4 1.2–4.5
n.a. n.a.
a
All fluxes given are for diffusive flux to the atmosphere. n.a. indicates that insufficient data were given to permit calculation of flux.
input to rivers, assumptions about the fraction of inorganic nitrogen species removed by nitrification or denitrification, and the fractional ‘yield’ of nitrous oxide production during these processes indicate that nitrous oxide fluxes to the atmosphere from estuaries is about 0.06–0.34 Tg N2O y1. Carbon Dioxide (CO2) and Oxygen (O2)
Estuaries are typically heterotrophic systems, which means that the amount of organic matter respired within the estuary exceeds the amount of organic matter fixed by primary producers (phytoplankton and macrophytes). Since production of carbon dioxide then exceeds biological removal of carbon dioxide, it follows that estuaries are likely to be sources of the gas to the atmosphere. At the same time, since oxidation of organic matter to CO2 requires oxygen, the heterotrophic nature of estuaries suggests that they represent sinks for atmospheric oxygen. In many estuaries, primary productivity is severely limited by the amount of light that penetrates into the water due to high particulate loadings in the water. In addition, large amounts of organic matter may be supplied to the estuary by runoff from agricultural and forested land, from ground water, from
sewage effluent, and from organic matter in the river itself. There are many reports of estuarine systems with oxygen saturations below 1 (undersaturated with respect to the atmosphere), but few studies in which oxygen flux to the estuary has been reported. However, estuaries are often dramatically supersaturated with respect to saturation with CO2, especially at low salinities, and a number of workers have reported estimates of carbon dioxide flux from these systems (Table 3).
Dimethylsulfide (DMS)
DMS is an atmospheric trace gas that plays important roles in tropospheric chemistry and climate regulation. In the estuarine environment, DMS is produced primarily from the breakdown of the phytoplankton osmoregulator 3-(dimethylsulfonium)-propionate (DMSP). DMS concentrations reported in estuaries are generally supersaturated, ranging from 0.5 to 22 nM, and increase with increasing salinity. DMS levels in the water column represent a balance between tightly coupled production from DMSP and microbial consumption. Only 10% of DMS produced from DMSP in the estuarine water column is believed to escape to the atmosphere from estuarine surface water, since the biological turnover of DMS (turnover time of 3–7 days) is approximately 10 times faster than DMS exchange across the air/water interface. A large part of the estuarine DMS flux to the atmosphere may occur over short time periods on the order of weeks, corresponding to phytoplankton blooms. The DMS flux for an estuary in Florida, USA, was estimated to be on the order of o1 nmoles m2 h1. Insufficient
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6
GAS EXCHANGE IN ESTUARIES
data are available to determine reliable global DMS air/water exchanges from estuaries. Hydrogen (H2)
Hydrogen is an important intermediate in many microbial catabolic reactions, and the efficiency of hydrogen transfer among microbial organisms within an environment helps determine the pathways of organic matter decomposition. Hydrogen is generally supersaturated in the surface waters of the few estuaries that have been analyzed for dissolved hydrogen with R-values of 1.5–67. Hydrogen flux to the atmosphere from estuaries has been reported to be on the order of 0.06–0.27 nmol m2 h1. The contribution of estuaries to the global atmospheric H2 flux cannot be determined from the few available data. Carbon Monoxide (CO)
Carbon monoxide in surface waters is produced primarily from the photo-oxidation of dissolved organic matter by UV radiation. Since estuarine waters are characterized by high dissolved organic carbon levels, the surface waters of estuaries are highly supersaturated and are a strong source of CO to the atmosphere. Reported R-values for CO range from approximately 10 to 410 000. Because of the highly variable distributions of dissolved CO within surface waters (primarily as the result of the highly variable production of CO), it is impossible to derive a meaningful value for CO emissions to the atmosphere from estuaries. Carbonyl Sulfide (OCS)
Carbonyl sulfide makes up approximately 80% of the total sulfur content of the atmosphere and is the major source of stratospheric aerosols. Carbonyl sulfide is produced within surface waters by photolysis of dissolved organosulfur compounds. Therefore, surface water OCS levels within estuaries exhibit a strong diel trend. Carbonyl sulfide is also added to the water column by diffusion from anoxic sediments, where its production appears to be coupled to microbial sulfate reduction. Diffusion of OCS from the sediment to the water column accounts for B75% of the OCS supplied to the water column and is responsible for the higher OCS concentrations in estuaries relative to the open ocean. While supersaturations of OCS are observed throughout estuarine surface waters, no trends with salinity have been observed. Atmospheric OCS fluxes to the atmosphere from Chesapeake Bay have been reported to range from 10.4 to 56.2 nmol m2 h1. These areal fluxes are over 50 times greater than those determined for the open ocean.
Elemental Mercury (Hg0)
Elemental mercury is produced in estuarine environments by biologically mediated processes. Both algae and bacteria are able to convert dissolved inorganic mercury to volatile forms, which include organic species (monomethyl- and dimethyl-mercury) and Hg0. Under suboxic conditions, elemental mercury also may be the thermodynamically stable form of the metal. In the Scheldt River estuary, Hg0 correlated well with phytoplankton pigments, suggesting that phytoplankton were the dominant factors, at least in that system. Factors that may affect elemental mercury concentrations include the type of phytoplankton present, photo-catalytic reduction of ionic Hg in surface waters, the extent of bacterial activity that removes oxygen from the estuary, and removal of mercury by particulate scavenging and sulfide precipitation. Fluxes of elemental mercury to the atmosphere have been estimated for the Pettaquamscutt estuary in Rhode Island, USA, and for the Scheldt, and range from 4.2–29 pmol m2 h1, although the values are strongly dependent on the model used to estimate gas exchange coefficients. Volatile Organic Compounds (VOCs)
In addition to gases produced naturally in the environment, estuaries tend to be enriched in byproducts of industry and other human activity. A few studies have investigated volatile organic pollutants such as chlorinated hydrocarbons (chloroform, tetrachloromethane, 1,1-dichloroethane, 1,2-dichloroethane, 1,1,1-trichloroethane, trichloroethylene and tetrachloroethylene) and monocyclic aromatic hydrocarbons (benzene, toluene, ethylbenzene, oxylene and m- and p-xylene). Concentrations of VOCs are controlled primarily by the location of the sources, dilution of river water with clean marine water within the estuary, gas exchange, and in some cases, adsorption onto suspended or settling solids. In some cases (for example, chloroform) there also may be natural biotic sources of the gas. Volatilization to the atmosphere can be an important ‘cleansing’ mechanism for the estuary system. Since the only estuaries studied to date are heavily impacted by human activity (the Elbe and the Scheldt), it is not possible to make generalizations about the importance of these systems on a global scale.
Conclusions Many estuaries are supersaturated with a variety of gases, making them locally, and occasionally
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GAS EXCHANGE IN ESTUARIES
regionally, important sources to the atmosphere. However, estuarine systems are also highly variable in the amount of gases they contain. Since most estuaries studied to date are in Europe or the North American continent, more data are needed before global budgets can be reliably prepared.
Glossary Air–water interface The boundary between the gaseous phase (the atmosphere) and the liquid phase (the water). Catabolic Biochemical process resulting in breakdown of organic molecules into smaller molecules yielding energy. Denitrification Reduction of nitrate via nitrite to gaseous endproducts (nitrous oxide and dinitrogen gas). Ebullition Gas transport by bubbles, usually from sediments. Estuary Semi-enclosed coastal body of water with free connection to the open sea and within which sea water is measurably diluted with fresh water derived from land drainage. Gas solubility The amount of gas that will dissolve in a liquid when the liquid is in equilibrium with the overlying gas phase. Henry’s Law Constant Proportionality constant relating the vapor pressure of a solute to its mole fraction in solution. Liquid laminar thickness The thickness of a layer at the air/water interface where transport of a dissolved species is controlled by molecular (rather than turbulent) diffusion. Molecular diffusivity (D) The molecular diffusion coefficient. Nitrification Oxidation of ammonium to nitrite and nitrate. Practical salinity scale A dimensionless scale for salinity. Redfield stoichiometry Redfield and colleagues noted that organisms in the sea consistently removed nutrient elements from the water in a fixed ratio (C : N : P=106 : 16 : 1). Subsequent workers have found that nutrient concentrations in the sea typically are present in those same ratios. Transfer coefficient The rate constant which determines the rate of transfer of gas from liquid to gas phase.
See also Air–Sea Gas Exchange. Air–Sea Transfer: Dimethyl Sulfide, COS, CS2, NH4, Non-Methane
7
Hydrocarbons, Organo-Halogens. Air–Sea Transfer: N2O, NO, CH4, CO.
Further Reading Bange HW, Rapsomanikis S, and Andreae MO (1996) Nitrous oxide in coastal waters. Global Biogeochemical Cycles 10: 197--207. Bange HW, Dahlke S, Ramesh R, Meyer-Reil L-A, Rapsomanikis S, and Andreae MO (1998) Seasonal study of methane and nitrous oxide in the coastal waters of the southern Baltic Sea. Estuarine, Coastal and Shelf Science 47: 807--817. Barnes J and Owens NJP (1998) Denitrification and nitrous oxide concentrations in the Humber estuary, UK and adjacent coastal zones. Marine Pollution Bulletin 37: 247--260. Cai W-J and Wang Y (1998) The chemistry, fluxes and sources of carbon dioxide in the estuarine waters of the Satilla and Altamaha Rivers, Georgia. Limnology and Oceanography 43: 657--668. de Angelis MA and Lilley MD (1987) Methane in surface waters of Oregon estuaries and rivers. Limnology and Oceanography 32: 716--722. de Wilde HPJ and de Bie MJM (2000) Nitrous oxide in the Schelde estuary: production by nitrification and emission to the atmosphere. Marine Chemistry 69: 203--216. Elkins JW, Wofsy SC, McElroy MB, Kolb CE, and Kaplan WA (1978) Aquatic sources and sinks for nitrous oxide. Nature 275: 602--606. Frankignoulle M, Abril G, Borges A, et al. (1998) Carbon dioxide emission from European estuaries. Science 282: 434--436. Frost T and Upstill-Goddard RC (1999) Air–sea exchange into the millenium: progress and uncertainties. Oceanography and Marine Biology: An Annual Review 37: 1--45. Law CS, Rees AP, and Owens NJP (1992) Nitrous oxide: estuarine sources and atmospheric flux. Estuarine, Coastal and Shelf Science 35: 301--314. Liss PS and Merlivat L (1986) Air–sea gas exchange rates: introduction and synthesis. In: Baut-Menard P (ed.) The Role of Air–Sea Exchange in Geochemical Cycling, pp. 113--127. Dordrecht: Riedel. Muller FLL, Balls PW, and Tranter M (1995) Processes controlling chemical distributions in the Firth of Clyde (Scotland). Oceanologica Acta 18: 493--509. Raymond PA, Caraco NF, and Cole JJ (1997) Carbon dioxide concentration and atmospheric flux in the Hudson River. Estuaries 20: 381--390. Robinson AD, Nedwell DB, Harrison RM, and Ogilvie BG (1998) Hypernutrified estuaries as sources of N2O emission to the atmosphere: the estuary of the River Colne, Essex, UK. Marine Ecology Progress Series 164: 59--71. Sansone FJ, Rust TM, and Smith SV (1998) Methane distribution and cycling in Tomales Bay, California. Estuaries 21: 66--77.
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GAS EXCHANGE IN ESTUARIES
Sansone FJ, Holmes ME, and Popp BN (1999) Methane stable isotopic ratios and concentrations as indicators of methane dynamics in estuaries. Global Biogeochemical Cycles 463--474. Scranton MI, Crill P, de Angelis MA, Donaghay PL, and Sieburth JM (1993) The importance of episodic events
in controlling the flux of methane from an anoxic basin. Global Biogeochemical Cycles 7: 491--507. Seitzinger SP and Kroeze C (1998) Global distribution of nitrous oxide production and N input in freshwater and coastal marine ecosystems. Global Biogeochemical Cycles 12: 93--113.
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GELATINOUS ZOOPLANKTON L. P. Madin and G. R. Harbison, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1120–1130, & 2001, Elsevier Ltd.
Introduction
tissues, which may also deter predation. Large size also permits commensal crustaceans to live on or in the body. Thus, as we look at the diversity of gelatinous zooplankton, we should keep in mind the forces that have led to their remarkable convergence. It is impossible to deal in a short article with the entire range of phyla that have gelatinous representatives, so some of the major groups will be highlighted.
Gelatinous zooplankton comprise a diverse group of organisms with jellylike tissues that contain a high percentage of water. They have representatives from practically all the major, and many of the minor phyla, ranging from protists to chordates. The fact that so many unrelated groups of animals have independently evolved similar body plans suggests that gelatinous organisms reflect the nature of the open ocean environment better than any other group. Whether as predators or grazing herbivores, they seem particularly well adapted to life in the oligotrophic regions of the world oceans, where their diversity and abundance relative to crustacean zooplankton is often greatest. The gelatinous body plan has evolved in a world where physical parameters are relatively constant but food resources are sparse or unpredictable. Gelatinous zooplankton exhibit many common adaptations to this habitat.
Species of polycystine radiolarians form large gelatinous colonies up to several meters in length. Thousands of individual protists are embedded in a common gelatinous matrix from which their pseudopodia extend into the water. The combined efforts of individuals in the colony enable relatively large plankton (such as copepods) to be captured and ingested. In addition to the protistan members of the colony, the matrix also contains symbiotic dinoflagellates (zooxanthellae) that grow on the metabolites of the radiolarians. In turn, the radiolarians digest the zooxanthellae, so that these colonies are planktonic homologues of coral reefs.
•
Medusae
•
•
• •
Transparent tissues provide concealment in the upper layers of the ocean, an environment without physical cover. Transparency is less common below the photic zone. The high water content of gelatinous tissues gives the organisms a density very close to that of sea water. The resulting neutral buoyancy decreases the energy required to maintain depth, but may actually require more energy overall, because of drag. The environment lacks physical barriers, strong turbulence, and current shears, so that gelatinous bodies do not need great structural strength. However, fragility makes many species difficult to sample or handle, and excludes most from more energetic coastal environments. High water content and noncellular jelly permit rapid growth and large body sizes, which can act as, or produce, large surfaces for the collection of food. Relatively large size makes gelatinous animals too big to be attacked by some predators, while their high water content reduces the food value of their
Taxonomic Groups Radiolaria
The phylum Cnidaria has many gelatinous representatives, comprising various groups of medusae and the strictly oceanic siphonophores (see below). What are commonly called jellyfish are medusae belonging to three Classes of the Cnidaria — the Hydrozoa, the Scyphozoa, and the Cubozoa. Since the morphology and life history of all three groups is broadly similar, it is practical to treat them together here. There are perhaps 1000 species of hydro- and scyphomedusae, probably with more to be discovered, especially in deep or polar waters. All are carnivorous, capturing prey with specialized stinging cells, called nematocysts. A wide variety of prey is eaten by different medusae, ranging from larval forms and small crustaceans to other gelatinous animals and large fish. Many epipelagic medusae also harbor zooxanthellae, and presumably they share their resources in the same way as the polycystine radiolarians. Many of these medusae are part of a life history that alternates between a sessile, benthic, asexually reproducing polyp and a sexually
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9
10
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reproducing and dispersing planktonic medusa. However, many oceanic species have lost the polyp stage and evolved instead a variety of sexual and asexual reproductive mechanisms that do not require a benthic habitat. There are several classification schemes for Cnidarians; the group names given here are common usage, but these vary in different taxonomies. Anthomedusae This order of hydromedusae includes small species ranging in size from less than 1 mm to several centimeters. The umbrella is usually shaped like a tall bell, and gonads are almost always found on the sides of the central stomach. There are four radial canals connecting the stomach to a marginal ring canal. Tentacles occur in varying numbers around the umbrella margin and sometimes around the mouth. Anthomedusae alternate with polyp forms, but some also bud medusae directly (Figure 1A). Leptomedusae These medusae are generally flatter than a hemisphere. They usually have four radial canals, but sometimes eight or more, or canals that are branched. Gonads are located on the radial canals, and there may be various sense organs on the margin. The stomach is sometimes flat, and sometimes mounted on a peduncle that can be quite long. There are tentacles around the margin but not the mouth. Leptomedusae also alternate with hydroids, but some species produce new medusae by budding or fission (Figure 1B). Limnomedusae Both high and low umbrella shapes are found in this order. There are usually four radial canals, sometimes branched. Gonads are either on the stomach or on the radial canals, and there is alternation of generations. Species of limnomedusae live in brackish, fresh (one species), or marine environments. Trachymedusae These medusae in the order Trachylina do not alternate generations but develop young medusae directly from planula larvae, or by asexual budding. The umbrella is often high, with stiff mesoglea and well-developed muscle fibers. Most have eight unbranched radial canals and gonads located on them. Many trachymedusae live in deep water and are heavily pigmented (Figure 1D). Narcomedusae Also in the Trachylina, narcomedusae have direct development from planulae, with a larval stage that is often parasitic on other medusae. There are no radial canals, but the flat central stomach is very wide and, in some genera,
extends into radial stomach pouches. Tentacles are solid and stiff, and often extend aborally. Narcomedusae are common in epipelagic and mesopelagic environments; some are strong vertical migrators (Figure 1C). Coronatae This order of scyphomedusae includes mainly deep-water species. The umbrella is divided into a high central part and a thinner marginal part by a coronal groove. The margin of the bell is divided into lappets; sense organs and solid tentacles arise from the cleft between lappets. The mouth has simple lips and the gastrovascular cavity is often deeply pigmented (Figure 1G). Semaeostomae The familiar large jellyfish are mainly in this order of the Scyphozoa. The umbrella margin is divided into lappets, and bears sense organs and hollow tentacles. There is no coronal groove around the umbrella. The mouth opening is surrounded by four long oral arms, often frilled. Gonads are in folds of the subumbrella (Figure 1E). Rhizostomae Medusae in this order of the Scyphozoa are mainly coastal species and can attain large size. They lack tentacles for prey capture, and instead ingest small particles carried into numerous small mouth openings by water currents. Some species in tropical waters host intracellular symbiotic algae. Cubomedusae Medusae in the class Cubozoa also alternate with a benthic polyp form, although details of their life cycles are poorly known. Cubomedusae can be quite large, and have the most virulently toxic nematocysts of any Cnidarians. Some species are responsible for human fatalities. Cubomedusae are also unusual in possessing complex, image-forming eyes, which are not as well developed in other medusae (Figure 1F). Siphonophores
The Order Siphonophora comprises a large and diverse group of predatory Cnidarians in the Class Hydrozoa. Their complex life cycles and colonial morphology are very different from the relatively simple hydromedusae and it is practical to consider the siphonophores as a separate group. The colonial, or polygastric, phase of the life cycle is the largest and most familiar. In this stage, siphonophores consist of an assemblage of medusoid and polypoid zooids, which are budded asexually from a founding larval polyp. The colony may include a gas float, nectophores or swimming bells, and a series of stem
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Figure 1 Medusae. (A) Pandea conica, an anthomedusa about 2 cm high. (B) Aequorea macrodactyla, a leptomedusa about 10 cm in diameter. (C) Cunina globosa, a narcomedusa about 5 cm in diameter. (D) Benthocodon hyalinus, a trachymedusa about 3 cm in diameter. (E) Cyanea capillata, a semaeostome scyphomedusa which can attain 1 m in diameter. (F) Carybdea alata, a cubomedusa up to 15 cm high. (G) Atolla wyvillei, a coronate scyphomedusa up to 25 cm diameter. (All photographs by L. P. Madin.)
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groups made up of feeding polyps and tentacles. In some siphonophores the stem groups break off as dispersal and sexually reproductive stages called eudoxids. The colony can be thought of as an overgrown, polymorphic juvenile stage that eventually bears the sexually reproductive adults. These are small medusoid zooids called gonophores, which produce gametes. Siphonophores range in size from a few millimeters to over 30 m in length, and occur throughout the water column. All are predators on other small zooplankton, and many genera are known to be luminescent. The colonies are fragile, and usually break up into their various units when collected in plankton nets. For this reason, much of the taxonomy is based on the morphology of the pieces, principally nectophores, and the appearance of the intact colonial stage is not always known. The Order Siphonophora is divided into three suborders and 15 families. Cystonectae This suborder includes siphonophores that possess a float but no swimming bells, so they are at the mercy of ocean currents. The Portuguese manof-war is the most familiar example. It has a float so large that the animal rests on the surface, but most cystonect species have smaller floats and are wholly submerged. Cystonects have virulent nematocysts and capture large, soft-bodied prey such as fish and squids (Figure 2A). Physonectae These siphonophores have more complex colonies, comprising a small apical float, numerous swimming bells that form a nectosome, and a stem containing several groups of gastrozooids, tentacles, bracts, etc. The stem typically contracts when the animal is swimming, and then relaxes so that the stem and tentacles extend to maximum length for fishing. This group is a major contributor to the deep scattering layer in many regions of the ocean. The largest siphonophores (the Apolemiidae, over 30 m long) are found in this group. Physonects prey mainly on small zooplankton, and many species are strong swimmers and vertical migrators. (Figure 2C). Calycophorae In this group, which contains the largest number of species, the float is absent and the nectophores are reduced, usually to two. A sequence of stem groups is budded and breaks free as eudoxids. Calycophorans are the most diverse, widely distributed, and abundant siphonophores. They catch small zooplankton and, when feeding, their tentacles form complex three-dimensional structures in the water, reminiscent of spider webs (Figure 2B).
Ctenophores
Ctenophores are exclusively marine gelatinous animals all but a few of which are holoplanktonic. Although they superficially resemble the Cnidaria, morphological and molecular studies indicate that Cnidarians and ctenophores are not closely related. Ctenophores are predators that use tentacles equipped with ‘glue’ cells or colloblasts to capture prey. The name ‘ctenophore’ is Greek for ‘comb bearer,’ referring to the comb-like plates of fused cilia that are used for propulsion. All ctenophores initially have eight meridional rows of comb plates, although in some groups these are lost or reduced during development. The vast majority of ctenophore species fall into six orders. Cydippida This group contains many species with paired tentacles that exit the body through tentacle sheaths. Species in the family Pleurobrachiidae catch prey ranging from small crustaceans to fish, while members of the Lampeidae feed mainly on large gelatinous animals like salps. The members of one species of cydippid, Haeckelia rubra, eat medusae, and retain the nematocysts of their prey (‘kleptocnidae’) for defensive use in their own tentacles. Before this behavior was known, these nematocysts were considered strong evidence for a close relationship between Cnidarians and ctenophores (Figure 2F). Platyctenida This group is primarily benthic and is distributed widely from the Arctic to the Antarctic. Members of the family Ctenoplanidae have comb rows as adults and are found in the plankton in the Indo-Pacific; all other species in the order lose their comb rows as adults, and live primarily as creeping benthic organisms. Platyctenes have functional tentacles that capture prey. Thalassocalycida This order contains a single species, Thalassocalyce inconstans, which lives in the midwater zone. It superficially resembles a medusa in overall shape, but can easily be distinguished by its eight comb rows and paired tentacles. Lobata Members of this order all have oral lobes and auricles, specialized structures that are used in feeding. Most lobates move through the water with their oral lobes widely spread to form a sort of basket. Small prey, such as crustaceans, are trapped on the mucus-covered oral lobes and tentilla, which stream over the body or extend onto the oral lobes. Ctenophores in the family Ocyropsidae lack
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Figure 2 Siphonophores and Ctenophores. (A) Rhizophysa filiformis, a cystonect siphonophore up to 2 m long. (B) Sulculeolaria sp. a calycophoran siphonophore up to 1 m long. (C) Physophora hydrostatica, a physonect siphonophore about 10 cm high. (D) Ocyropsis maculata, a lobate ctenophore about 8 cm in diameter. (E) Beroe cucumis, a beroid ctenophore up to 25 cm long. (F) Mertensia ovum, a cydippid ctenophore about 4 cm high. (Photographs A, C, D, F by G. R. Harbison; B, E by L. P. Madin.)
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tentacles, and capture prey by enclosing them in their muscular oral lobes (Figure 2D). Cestida These ctenophores are shaped like long, flat belts. They appear to be related to the Lobata, but lack oral lobes and auricles. There are only two genera (Cestum and Velamen) in one family (Cestidae). The comb rows extend along the aboral edge of the ribbonlike body, propelling the animal with the oral edge forward. Small prey are captured by the fine branches of the tentacles that lie over the flat sides of the body. Cestids are characteristic of oceanic, epipelagic environments. Beroida Beroids lack tentacles altogether. Their large stomodaeum occupies most of the space in the body. All beroids are predators on other ctenophores, and occasionally salps. They capture prey by engulfing them, and can bite off pieces of the prey with specialized macrocilia located immediately behind the mouth (Figure 2E).
others of which (Pseudothecosomata) have gelatinous shells and tissues. There are over 30 species of euthecosome pteropods, in two families, the Limacinidae and the Cavoliniidae. Thecosome pteropods feed by collecting particulate food on the surface of a mucous web or bubble, produced by mucus glands on the wings, and held above the neutrally buoyant and motionless animal. The mucus is periodically retrieved and ingested along with adhering particles, then replaced by a newly secreted web. Some cavoliniids have brightly colored mantle appendages that may aid in maintaining neutral buoyancy or serve as warning devices to predators. When disturbed, animals lose their neutral buoyancy and rapidly sink (Figure 3A). The Pseudothecosomata includes three families, the Peraclididae (one genus), the Cymbuliidae (three genera), and the Desmopteridae (one genus). Pseudothecosomes are larger than euthecosomes, and their mucous webs are correspondingly larger, reaching over a meter across in Gleba cordata (Figure 3C).
Heteropods
The Phylum Mollusca contains many gelatinous representatives, and the gelatinous body plan has apparently arisen independently in several groups. The Heteropoda is a superfamily of prosobranch gastropods that includes the families Atlantidae, Carinariidae, and Pterotracheidae. Heteropods are visual predators with well-developed eyes and a long proboscis containing a radula. Atlantid heteropods have thin, flattened shells into which they can completely withdraw their bodies. They feed on small crustaceans and other mollusks. The family Carinariidae includes eight species in three genera, Carinaria, Pterosoma, and Cardiapoda. These heteropods have a greatly reduced shell, enclosing only a small fraction of the body. Carinariids feed primarily on other gelatinous organisms, such as salps, doliolids and chaetognaths. In the family Pterotracheidae, with two genera — Pterotrachea (four species) and Firoloida (one species) — the shell is completely absent (Figure 3D). Pteropods
This molluskan group comprises two orders in the gastropod subclass Opisthobranchia. The foot in pteropods is modified into two wingshaped paddles responsible for swimming; their fluttering gives rise to the common name for pteropods, sea butterflies. Thecosomata This group contains the shelled pteropods, some of which (Euthocosomata) have calcareous shells and are not truly gelatinous, and
Gymnosomata Members of this order are poorly known, largely because they have no shells and contract into shapeless masses when preserved. Most species live in the deep sea, and only a few of the approximately 50 species have been observed alive. Gymnosomes appear to be highly specialized predators on particular species of thecosome pteropods. The order is divided into two suborders, the Gymnosomata and the Gymnoptera. The four families of the Gymnosomata include the Pneumodermatidae (seven genera and 22 species) with sucker-bearing arms similar to cephalopod tentacles; the Notobranchaeidae (one genus and eight species), with suckerless feeding arms called buccal cones; the Clionidae (eight genera and 16 species); and the Cliopsidae (two genera and three species) (Figure 3B). There are two families in the Gymnoptera, the Hydromylidae (one genus, one species) and the Laginiopsidae (one genus, one species). These groups are very different from each other and from other gymnosome pteropods, and some may not actually belong in the Order Gymnosomata. Cephalopods
Although many cephalopods are active, muscular swimmers, there are several gelatinous and/or transparent genera. The family Cranchiidae is composed entirely of gelatinous species, including the genera Taonius, Megalocranchia, and Teuthowenia. These relatively large, slow-moving squids probably
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Figure 3 Mollusks and Polychaete. (A) Cuvierina columnella, a euthecosome pteropod about 1 cm high. (B) Clione limacina, a gymnosome pteropod about 2 cm high. (C) Corolla spectabilis, a pseudothecosome pteropod about 10 cm in diameter, with mucous web in background. (D) Carinaria sp. a heteropod about 10 cm long. (E) Teuthowenia megalops, a cranchiid cephalopod about 10 cm long. (F) Alciopid polychaete worm, up to 1 m long. (Photographs A, B, C, E, F by G. R. Harbison; D by L. P. Madin.)
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capture prey through stealth rather than active pursuit. Vitreledonelliid octopods are also gelatinous (Figure 3E). Polychaete Worms
Two major groups of planktonic polychaetes are gelatinous, the Alciopidae and the Tomopteridae. Both are in the order Phyllodocida, although they are probably not closely related. Alciopids are characterized by well-developed eyes with lenses. Many have ink glands along the sides of their bodies, which may function analogously to the ink glands of cephalopods. Their habits are poorly known, but they may feed on gelatinous prey. Alciopids may attain lengths of nearly a meter. Tomopterids do not have well-developed eyes, but probably use chemoreception to locate prey. Some release luminous secretions from glands along their body when disturbed. Deep-sea tomopterids may be 25 cm long, but most shallow species are much smaller. (Figure 3F). Crustaceans
Although arthropods cannot really be considered gelatinous because of their exoskeletons, there are some examples of very transparent bodies, presumably also an adaptation for concealment. The most notable examples are species of the hyperiid amphipods Cystisoma and Phronima. Species of Cystisoma are large and transparent, and the enormous retinas of the compound eye are lightly tinted. Although the retinas of species of Phronima are darkly pigmented, the rest of the body is transparent. It is likely that the transparency of these species is a form of protective coloration, since they live on transparent gelatinous hosts. Holothurians
Although the majority of holuthurian species are rather sedentary benthic deposit feeders, there are several deep-sea genera of swimming or drifting holothurians with gelatinous bodies. Species in the genera Peniagone and Enypniastes feed on bottom deposits, but can swim up into the water column. The genus Pelagothuria appears to be wholly pelagic, with a morphology that suggests it collects and feeds on sinking particulate matter. Few pelagic holuthurians have been observed alive and little is known of their life history or behavior. Pelagic Tunicates
The subphylum Urochordata includes two classes of pelagic tunicates, the Thaliacea and the Larvacea or
Appendicularia. Thaliaceans (including the orders Pyrosomida, Doliolida, and Salpida) are relatively large animals with more or less barrel-shaped bodies. They pump a current of water through their bodies and strain phytoplankton and other small particulates from it with a filter made of mucous strands. The same current provides jet propulsion. Thaliaceans have complex life cycles with alternating generations and multiple zooid types. The class Larvacea comprises a single order of small organisms that filter food particles using an external mucous structure called a house. Both Thaliaceans and Larvaceans are widely distributed in the oceans, and are sometimes extremely abundant. Pyrosomida Pyrosomes form colonies made up of numerous small ascidian-like zooids embedded in a stiff gelatinous matrix or tunic. The colony is tubular, with a single terminal opening. Water is pumped by ciliary action through each zooid, and suspended food particles are retained on the branchial basket within the body. The excurrent water from each zooid passes into the lumen of the colony, forming a single exhalent current that provides jet propulsion. Most pyrosome colonies are a few centimeters to a meter in length, but colonies of at least one species can attain lengths of 20 m. Doliolida This order of the Thaliacea comprises six genera and 23 species of small (2–10 mm), barrel-shaped animals with circumferential muscle bands. The filter feeding mechanism is similar to that of pyrosomes, with currents generated by ciliary beating passing through a mucous net supported on the branchial basket. The life cycle involves five asexual stages and one sexual stage, several of which occur together as parts of large colonies of thousands of zooids. These colonies may attain lengths over 1 m, but are fragile and are rarely collected intact. In most genera of doliolids, the life cycle begins with a sexually produced larva, which becomes the oozooid stage. This stage feeds initially, but then begins budding off the trophozooid and phorozooid stages, thus forming the colony. During this process the oozooid loses its branchial basket and gut, and transforms into the ‘old nurse’ stage, whose function is to swim by jet propulsion and pull the attached colony along behind it. Contractions of the body muscles produce short exhalent pulses that move the colony rapidly. The trophozooids in the colony filter-feed to support themselves and the nurse. The phorozooids grow attached to the colony, but then break free to lead independent lives and produce asexually a
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small group of gonozooids. These eventually break free from the phorozooid, and become the sexually reproducing stages (hermaphrodites?) that produce the larvae and begin the whole cycle again (Figure 4B). Salpida This order (with 12 genera and about 40 species) is of larger filter-feeding animals, also with
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circumferential muscle bands. The salps alternate between two forms, an asexually budding solitary (oozooid) stage and a sexually reproducing aggregate (blastozooid) stage. The aggregate salps usually remain connected together in chains or whorls of various types. Swimming is by jet propulsion, produced by a pulsed water current generated by rhythmic contraction of body muscles.
Figure 4 Pelagic Tunicates. (A) Megalocercus huxleyi, a larvacean of about 5 mm body length, house length about 4 cm. (B) Dolioletta gegenbauri, portion of a colony showing gastrozooids and phorozooids, individuals 2–5 mm long, colonies up to 1 m. (C) Salpa maxima, solitary generation salp, up to 25 cm long. (D) Salpa maxima, chain of aggregate generation salps; orange dots are guts of salps; individuals are to 15 cm, chains up to 10 m long. (E) Pegea socia, aggregate generation salp with attached embryo of solitary generation; aggregate 7 cm, embryo about 1 cm. (F) Traustedtia multitentaculata, solitary generation salp with appendages of uncertain function, about 3 cm long. (All photographs by L. P. Madin.)
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Food particles are strained from the water passing through the body cavity by a mucous filter, which is continuously secreted and ingested. The individual animals range in size from 5 to over 100 mm, and chains can be several meters long (Figures 4C–F). Larvacea This class (also called Appendicularia) is divided into three families (with 15 genera and 70 species) of small (1–10 mm) animals consisting of a trunk and long, flat tail. Larveaceans are also filter feeders on small particulates but are unique among tunicates in the use of an external concentrating and filtering structure called the house. The house surrounds the animal, and contains a complex set of channels and filters made of mucous fibers and sheets. Water is pumped into the house by the oscillation of the larvacean’s tail; the exhalent stream provides slow jet propulsion in some species. Particles are sieved from the flow as it passes through the internal filter; they accumulate and are aspirated at intervals into the pharynx of the larvacean via a mucous tube. The complex house is formed as a mucous secretion on the body of the larvacean, produced by specialized secretory cells. It is inflated with sea water, pumped into it by action of the tail, until it attains its full size, with all the internal structures expanding in proportion. Houses eventually become clogged with particulates and fecal pellets, and are then jettisoned. The larvacean expands a new house (there may be several house rudiments on its body, awaiting expansion) and resumes filter feeding. The abandoned houses can be an important source of marine snow and serve as food for various planktonic scavengers (Figure 4A).
Ecology of Gelatinous Zooplankton Gelatinous zooplankton are found in all of the oceans of the world, from the tropics to polar regions. They also occur at all depths, and many of the largest and most delicate species have been collected in recent years from the mesopelagic and bathypelagic parts of the ocean. The absence of turbulence in the deep sea probably allows these species to attain such large sizes, but there are also robust species that thrive in surface and coastal environments. Examples include the Portuguese man-of-war (Physalia physalis), which lives at the air–water interface and can ride out hurricanes, and the ctenophore Mnemiopsis leidyi, which lives in estuaries with strong tidal currents and turbulence. In general, gelatinous organisms have been rather neglected by zooplankton ecologists, primarily because their delicacy makes them difficult to sample
and study. Most are damaged or destroyed in conventional plankton nets, and many deep-water siphonophores and ctenophores are too delicate to be captured intact even with the most gentle of techniques. Much recent progress in understanding their biology has been based on in situ methods of study using SCUBA diving, submersibles, or remote vehicles. These methods permit observation of undisturbed behavior and collection of intact living specimens. Advances in culture techniques and laboratory measurements have improved our understanding of energetics, reproduction, and life history of some species, but most remain only partially understood. Gelatinous animals occupy every trophic niche, ranging from primary producers (symbiotic colonial radiolaria) to grazers (pteropods and pelagic tunicates) and predators (medusae, siphonophores, and ctenophores). In all these niches, the gelatinous body plan confers advantages of size and low metabolic costs. In addition to attaining large sizes with relatively little food input, gelatinous organisms such as medusa and ctenophores are able to ‘de-grow’ when deprived of food. Metabolic rates remain unchanged, and the animal simply shrinks until higher food levels allow it to resume growth. This energetic flexibility is probably important to the success of gelatinous species in the oligotrophic open ocean and deep sea. Many species of medusae and siphonophores, for example, appear able to survive at low population densities spread over very large areas. In other cases the efficiency of their feeding, growth, and reproduction allows gelatinous species to outcompete other types of zooplankton and form dense populations over large areas, which can have considerable impact on ecosystems. A dramatic recent example was the accidental introduction of the ctenophore Mnemiopsis leidyi into the Black Sea from the eastern seaboard of the Americas. In the late 1980s, this ctenophore reproduced in prodigious quantities, and the resulting predation on zooplankton and larval fishes led to the collapse of pelagic fisheries in the Black Sea. These fisheries have to some extent recovered, but seasonal blooms of this ctenophore continue to occur in the Black Sea, just as they do on the eastern shores of the Americas. Reports of Mnemiopsis in the Caspian Sea suggest that the pattern may be repeated. Many other gelatinous species form dramatic seasonal blooms, such as the medusa Chrysaora quinquecirrha in the Chesapeake Bay, the salp Thalia democratica off Florida, Georgia, and Australia, and the medusa Pelagia noctiluca in the Mediterranean. In the Southern Ocean, immense populations of the salp Salpa thompsoni alternate with those of the
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Antarctic krill Euphausia superba. The formation of large aggregations through rapid reproduction appears to be a common strategy for taking advantage of favorable conditions. Dense populations are sometimes further concentrated by wind or current action, or are transported close to the coast from their normal habitats farther offshore. The combination of rapid growth and advection can cause the sudden appearance of swarms of medusae, ctenophores, or salps in coastal waters. Although these blooms may sometimes have serious or even catastrophic effects on other organisms, including fisheries or human activities, they are a natural part of the life histories of the species, and not events for which remedial action is needed, or even possible. Gelatinous zooplankton are normal components of virtually all planktonic ecosystems. They are among the most common and typical animals in the oceans, whose biology and ecological roles are now becoming better understood.
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See also Plankton. Zooplankton Sampling with Nets and Trawls.
Further Reading Bone Q (1998) The Biology of Pelagic Tunicates. Oxford: Oxford University Press. Harbison GR and Madin MP (1982) The Ctenophora. In: Parker SB (ed.) Synopsis and Classification of Living Organisms. New York: McGraw Hill. Lalli CM and Gilmer RW (1989) Pelagic Snails. Stanford, CA: Stanford University Press. Mackie GO, Pugh PR, and Purcell JE (1987) Siphonophore biology. Advances in Marine Biology 24: 98--263. Needler-Arai M (1996) A functional biology of Scyphozoa. London: Chapman and Hall. Wrobel D and Mills CE (1998) Pacific Coast Pelagic Invertebrates: A Guide to the Common Gelatinous Animals. Monterey: Sea Challengers and Monterey Bay Aquarium.
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GENERAL CIRCULATION MODELS G. R. Ierley, University of California San Diego, La Jolla, CA, USA
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Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1130–1134, & 2001, Elsevier Ltd.
Introduction A general circulation model (GCM) of the ocean is nothing more than that – a numerical model that represents the movement of water in the ocean. Models, and more particularly, numerical models, play an ever-increasing role in all areas of science; in geophysics broadly, and in oceanography specifically. It was perhaps less the early advent of supercomputers than the later appearance of powerful personal workstations (tens of megaflops and megabytes) that effected not only a visible revolution in the range of possible computations but also a more subtle, less often appreciated, revolution in the very nature of the questions that scientists ask, and the answers that result. The range of length scales and timescales in oceanography is considerable. Important dynamics, such as that which creates ‘salt fingers’ and hence influences the dynamically significant profile of density versus depth, takes place on centimeter scales, while the dominant features in the average circulation cascade all the way to basin scales of several thousand kilometers. Timescales for turbulent events, like waves breaking on shore, are small fractions of seconds, while at the opposite end, scientists have reliably identified patterns in the ocean with characteristic evolution times of order several decades.1 Most simply, a ‘model’ is no more than a mathematical description of a physical system. In the case of physical oceanography, that description includes the following elements:
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The momentum equation (F ¼ mv_ , but expressed in terms appropriate to a continuous medium), or often in its place a derivative form, the ‘vorticity equation’.
1 It is useful to distinguish extrinsic evolution times, which span geologic time, from intrinsic variation, which characterizes an isolated ocean and atmosphere, thus neglecting such secular influences as orbital variation, change in the earth’s rotation rate, variation in the solar constant, etc. Although not all causes of variability have been identified, it is possible that even documented evolution over thousands of years may reflect the latter, intrinsic, variability.
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• • •
An equation to express the principle of mass conservation. A heat equation, which describes the advection (carrying by the fluid) and diffusion of temperature. A similar ‘advection–diffusion’ equation for salinity. An equation of state, which relates the pressure to the density.
Although it may, after the fact, sound obvious, it took scientists many years to appreciate the full role of rotation – which enters Newton’s law through a latitude dependent Coriolis force – in generating the observed large scale circulation of ocean. Elements of these equations can become quite complicated. For example, momentum in the upper ocean is imparted by a complex, not yet fully understood, process of wind–wave interaction. One has to choose whether to represent the action of the wind kinematically, which means that the spatial and temporal variability of the wind must be given beforehand, or dynamically, in which case we must solve not only the set of equations above, but a similar set that describes the simultaneous evolution of the atmosphere. Clearly the dynamical case is the more ‘realistic’ of the two, but the price is a substantially more involved computation. It is issues such as these that force one to a choice that often pits understanding and intuition on the one hand against verisimilitude and complete dynamical consistency on the other. As the common aim of most large-scale models is to generate results of maximum realism, it usual that the models are frequently corrected, or ‘steered’, on the fly by extensive use of data assimilation. It is not merely a matter of slight refinement: no large-scale model (GCM) is yet sufficiently robust that it can be used for forecasting without considerable input of such observationally derived constraints; in the oceanic component, for example, the need to force model agreement at depth by continuously ‘relaxing’ the solution to the smoothed data of the Levitus atlas (a world-wide compendium of data from many sources, smoothed and interpolated onto a regular grid). How much of this fragility is because of explicit defects (faulty ‘subgrid scale modeling’ and explicit omissions in the physics e.g., neglect of wave breaking) and how much is due to discrete numerical implementation inconsistent with plausible continuous equations remains an open question. Their limitations notwithstanding, large numerical models are one vital means by which we grapple with questions about global warming and a host of other
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environmental issues that affect the way that both we and future generations will live.
instead we reach a tentative conclusion of not proven wrong (yet!). While it may stimulate conjecture, and have other worthy ends, the idea of introducing artificial parameters for the express purpose of manufacturing close agreement with reality is nonetheless formally antithetical to the paradigm of testing independent, quantitative predictions. Backward compatibility. As with confinement of form, the idea of compatibility achieves a purity in mathematics that is not to be expected of the physical sciences. Each new bit of mathematics must fit perfectly into the entire edifice of results already discovered (or invented, as you will). Commonly, though not without exception, in the physical sciences, newer theories are seen to encompass the older theories as special or limiting cases. We speak, for example, of the ‘classical limit’ as a means of recovering prerelativistic or prequantum results. Indeed, it is only in light of Einstein’s theory of special relativity that we can understand the limitations of Newton’s law, which we now understand more fully as not a law, but a limiting approximation. Oceanography has families of theories, each nesting one within the next, like a series of oceanographic Matryoshka dolls, the innermost of which is often the theory of ‘quasigeostrophy,’ which dates from the 1950s.
Models in Theory and Practice Historically, computers were initially so limited that the questions posed were often, in effect, slight extensions of preexisting analytic queries. To that extent, such studies continued to conform (at least in principle) to what we may identity as four basic building blocks of most theories, which, in aiming to describe physical reality become subject to constraints. Classical modes to which physical laws are found to conform generally2 include the following.
•
•
•
Expression in quantity, extent, and duration. The language that offers itself as encompassing all distinct physical conditions and all meaningful physical relations is fundamentally that of mathematics. As others before and since, the great mathematician Eugene Wigner too had a stab at explaining what, in an eponymous essay, he termed ‘the unreasonable effectiveness of mathematics in the physical sciences’. It remains a conundrum. Confinement of form. This attains its purest expression in the Platonic view that mathematical entities are not invented, but discovered, and as such have a prior, if not physical, existence. In extension to physical theories, it corresponds to the belief that there is some true, ultimate, equation association with any natural phenomenon. One important respect in which this idea must be tempered in application to fluids is the mathematical demonstration that a variety of different microscopic laws for interaction may all yield the same generic macroscopic law applicable to behavior at large space scales or timescales. If one’s aim is solely to understand the latter, then although for a given problem we might presume that there is indeed a precise, if complicated, microscopic law, one’s effort might more profitably be spent understanding the passage to the large-scale limit. Falsifiability. Implicit in the progress of science is the idea that one makes and then tests hypotheses. But unlike in mathematics, in the physical sciences we cannot show the hypothesis is correct, only that it is wrong. A hypothesis is never vindicated,
2 We speak gently here, since to insist that those four are always either necessary or sufficient would require that we introduce a theory about theories: a metatheory. But we have no notion of how rigorously to evaluate such ideas!
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•
The unavoidable adoption and resulting sensitivity of GCMs to ad hoc parameterizations (e.g., subgrid scale modeling) or necessity set them apart from the traditional pursuit of the scientific method. This distinction was (presciently) appreciated at least by the early 1960s and it changes, or ought to change, one’s view of such models as rigorous arbiters of precise truths. And yet while it is true that numerical experiments with GCMs are merely suggestive rather than truly predictive of future evolution of the ocean, the sheer lack of experimental data, to say nothing of the lack of a control, means that theoretical ideas are often assessed on the basis of their success in explaining strictly numerical experiments. As computers became more powerful it was natural to press for the most realistic model runs possible. And, because the growth in computing power was increasingly realized through distributed3 as well as mainframe (super), computing, the school of ‘kitchen sink’ models, which started as a specialized branch off the mainstream of oceanography – largely
3
Both virtually, through high speed and increasingly transparent networks, and physically, through desktop workstations of considerable power.
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limited in participation to those in close physical proximity to two or three central machines – became a powerful tributary in its own right: an autonomous discipline within oceanography, which naturally began to evolve its own criteria for relevance.
Limits on Numerical Models We spoke of two sources of error common to large numerical models: difficulties in numerical implementation, and poorly modeled or unrepresented physical processes. In this section we consider specific instances of each, starting with an abstract mathematical point of view. But note that if we could with the wave of hand dispense with these two issues (which one supposes are in principle tractable), the fact that we do not know the exact physical state of the ocean inevitably increases the uncertainty of the results. Moreover, even were we given that exact state at some instant in time, the intrinsic and spontaneous genesis of disorder in such a physical system must forever constrain our predictive power. There are two key mathematical features of the basic momentum (or vorticity) equation which bear comment: conservation and dissipation. The nonlinear term is fundamental to the initiation and sustenance of turbulence and by itself strictly conserves energy. (Other terms introduce explicit dissipation.) In addition, in two dimensions the term conserves ‘enstrophy’ (the square of the vorticity) and in three, both ‘circulation’ and ‘helicity’ (the dot-product of velocity and vorticity). While one might hope for all such conservation properties to be preserved in numerical implementations, some large models, often those based on curvilinear – as opposed to Cartesian – coo¨rdinates – manage only to conserve energy. Beyond the conserved quantities associated with the nonlinear term, which include the energy and, in general, the so-called ‘Noether invariants,’ there is a more subtle property associated with the exact (continuous) equation: its associated ‘multisymplectic geometry.’ Recent mathematical advances make it possible for a discrete numerical model to preserve such structure exactly, though as yet such improvements have not been incorporated into any working GCM. Is it quantitatively important that we do so when basic fluid processes, such as convection, are as yet only crudely modeled? Until the experiment is tried, no one can say. But it is pertinent to note that a similar (that is, ‘symplectic’) refinement is critical for a numerical solution of sufficient accuracy that one can decide whether planetary orbital motions are chaotic on astronomical timescales.4 Finally, strictly speaking, not only ambitious models, but even more confined ‘process’ models, rest
upon a not yet wholly secure foundation: it is still an open research question whether the basic equation of fluid mechanics (the Navier–Stokes equation) is itself ‘globally well-posed’ in three dimensions.5 (It is widely believed to be so, but belief comes cheap. A proof, however, is worth one million dollars(!) – one of seven prizes in a competition recently announced by the Clay Mathematics Institute.) Even overlooking such foundational matters on which mainly mathematicians would cavil about numerical models, the last three of the four principles above are often violated in more apparent ways in the application and development of present-day models. We illustrate this divergence from traditional norms with a few representative examples to emphasize the sometimes causal (not casual!) link between models and ‘reality.’ In delineating the borders of the known, the unknown, and the unknowable, it is important to discriminate between deduction and rationalization as competing processes for exploring and explaining those borders.
•
Although GCM simulations with a viscosity approaching that of water are at present inconceivable, at least as a thought experiment it is worth bearing in mind that those are the numerical results we would in principle compare against observation to assess a given model.6 Short of that, GCMs use various formulations that ostensibly mimic the dynamical effects of the unresolved scales of motions. In the simplest instance, this amounts to choosing a numerical viscosity several orders of magnitude larger than that of sea water. But often the value is dictated by purely heuristic numerical considerations: it is set at a threshold value, any decrease below which leads to rapid
4
On dissipation, a deep, though perhaps insufficiently appreciated, mathematical result is that the solution of a parabolic, dissipative system quickly collapses onto a finite-dimensional ‘attractor’. This is remarkable. If you think of assigning a point to every one of N molecules of water in the ocean, and tracking the velocity and position of each, the associated ‘phase space’ – just a record of that evolution – has dimension 6N. Because the momentum equation is derived on the basis that water is an infinitely divisible continuum, strictly we need to imagine that N approaches infinity. Nonetheless, even in that infinite-dimensional phase space, it remains true that the solution confines itself to only a finite, if quite large, portion. 5 Hadamard introduced the notion of ill-posedness of partial differential equations. A problem is well-posed when a solution exists, is unique, and depends continuously on the initial data. It is ill-posed when it fails to satisfy one or more of these criteria. 6 There is a curious division among physical oceanographers as to whether the large-scale flow we observe is, in the end, actually sensitive to the precise value of the viscosity of sea water. Predictably, there are two camps: yes and no.
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GENERAL CIRCULATION MODELS
numerical blowup. It cannot be said to be satisfactory feature that a basic parameter is set not by independent dynamical considerations but for stability reasons, and those pertaining solely to the discrete form of the equations. (The solution to the continuous equation would not blow up!) At times, not only the coefficient, but the actual form of the diffusive operator is adjusted. As above, usually the motivation is intrinsically numerical, so it is not surprising that a catalogue of the various choices shows some, for example, that create artificial sources (or sinks) of vorticity in the flow. Others, subject to the given boundary conditions, do not make mathematical sense in a region where the fluid depth tends to zero (like Atlantic City). Oceanographers, unfortunately, do not have the luxury of extensive laboratory measurements from which their dissipative parameterizations can be calibrated. A program of direct observation in particular regions where dissipation is thought to be significant is just getting under way. While such measurements will help reveal deficiencies in present formulations, the largest GCMs will probably rely on a solely heuristic approach for some time to come.
•
An important, but numerically unresolved, process in the ocean is that of convection, which typically occurs at small scales in, for example, localized regions of intense surface cooling. The overall thermal structure of the ocean is sensitive to this, a means by which ‘bottom water’ is formed; a cold, relatively less salty mass that constitutes the deep Atlantic, for example. Because the horizontal resolution is too coarse to encompass the sinking motions, various schemes have been devised to mimic that effect. It has been shown that one of the most common of these leads paradoxically to unacceptable physical (and mathematical) behavior as resolution is improved; it has no verifiable correspondence to a realizable physical process. The temptation with a model that has been extensively tuned to give plausible answers for other observables is to leave well enough alone. Unhappily, a model with one or more such elements whose limits are ill-defined or nonexistent must inevitably produce end results whose errors are typically an opaque mix of effects: some physical, some numerical, some mathematical. In such circumstances, the program of falsifiability of the physical components is apt to be fatally compromised.
The point is not that one should immediately dispense with all ad hoc parametrizations; excepting those that are simply mathematically ill-posed from
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the outset, theorists generally do not have better alternatives to suggest. But one should always bear in mind the degree to which numerical simulations are sensitive to these components, and seek independent ways in which to constrain their parameters, in isolated settings that test the limits of prediction against known measurements or, failing that, at least against fully converged, adequately resolved simulations of a local or regional character. If, within the acceptable parameter range identified, it is found that the original model no longer gives adequate large-scale predictions, then there are more basic problems to be addressed.
Summary From the numerical side, no computer improvements that can be seen on the horizon seem likely to make reasonably ambitious GCMs accessible to rigorous and extensive parametric and numerical exploration, a prerequisite to their complete understanding. From the mathematical side, it seems to be our fundamental ignorance about turbulence that most severely restricts the range of our grasp, leaving us with an often painfully narrow range of computations to which theoretical remarks can be significantly addressed. For these structural reasons, the gulf between theory and much numerical modeling will probably continue to widen for the foreseeable future, and thus there may grow to be—indeed some would say it already exists—a division akin to C.P. Snow’s ‘Two Cultures’. All the cautions about GCMs notwithstanding, they have become an integral part of the study of physical oceanography. With due regard for the novel capacities and limitations of numerical models, such scientific progress as we do make will more and more often hinge upon judicious computation.
See also Deep Convection. Double-Diffusive Convection. Forward Problem in Numerical Models. Wind Driven Circulation.
Further Reading The literature on ocean modeling is not yet productive of definitive treatises, in large measure because the field is yet young and rapidly evolving. Thus in lieu of textbooks or similar references, the reader is directed to the following series of articles. For some predictions on the perennially intriguing issue of what improvements in large-scale modeling may be
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driven by plausible increases in computing speed with massively parallel machines see Semtner A. Ocean and climate modeling Communications of the ACM 43 (43): 81–89. For a look back at the history of one of the single most influential models in physical oceanography, see A.J. Semtner’s Introduction to ‘A numerical method for the study of the circulation of the World Ocean’, which accompanies the reprinting of Kirk Bryan’s now classic 1969 article of the title indicated. This pair appears back-toback, beginning on page 149, in Journal of Computational Physics, (1997) 135 (2).
General readers may wish to consult the following succinct review, accessible to a broad audience: Semtner AJ (1995) Modeling ocean circulation Science 269 (5229): 1379–1385. Finally, for those readers desiring a more in depth appreciation of modeling issues and their implications for specific features of the large scale circulation, consult the careful review. McWilliams JC (1996) Modeling the oceanic general circulation. In: Lumley JL, Van Dyke M. Read HL (eds) Annual Review of Fluid Mechanics Vol. 28, pp. 215–248 Palo Alto, CA: Annual Reviews.
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GEOMAGNETIC POLARITY TIMESCALE C. G. Langereis and W. Krijgsman, University of Utrecht, Utrecht, The Netherlands Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1134–1141, & 2001, Elsevier Ltd.
maintained by convective fluid motion. At the surface of the Earth, the field can conveniently be described as a dipole field, which is equivalent to having a bar magnet at the center of the Earth. Such a dipole accounts for approximately 90% of the observed field; the remaining 10% derives from
Introduction Dating and time control are essential in all geoscientific disciplines, since they allows us to date, and hence correlate, rock sequences from widely different geographical localities and from different (marine and continental) realms. Moreover, accurate time control allows to understand rates of change and thus helps in determining the underlying processes and mechanisms that explain our observations. Biostratigraphy of different faunal and floral systems has been used since the 1840s as a powerful correlation tool giving the geological age of sedimentary rocks. Radiometric dating, originally applied mostly igneous rocks, has provided numerical ages; this method has become increasingly sophisticated and can now — in favorable environments — also be used on various isotopic decay systems in sediments. We are concerned with the application of magnetostratigraphy: the recording of the ancient geomagnetic field that reveals in lava piles and sedimentary sequences, intervals with different polarity. This polarity can either be normal, that is parallel to the present-day magnetic field (north-directed), or reversed (south-directed) (Figure 1). As a rule, it appears that these successive intervals of different polarity show an irregular thickness pattern, caused by the irregular duration of the successive periods of either normal or reversed polarity of the field. This produces a ‘bar code’ in the rock record that often is distinctive. Polarity intervals have a mean duration of some 300 000 y during the last 35 My, but large variations occur, from 20 000 y to several million years. If one can construct a calibrated ‘standard’ or a so-called ‘geomagnetic polarity timescale’ (GPTS), dated by radiometric methods and/or by orbital tuning, one can match the observed pattern with this standard and hence derive the age of the sediments. Magnetostratigraphy with correlation to the GPTS has become a standard tool in ocean sciences.
The Paleomagnetic Signal The Earth’s magnetic field is generated in the liquid outer core through a dynamo process that is
N
Normal
S
Reversed Figure 1 Schematic representation of the geomagnetic field of a geocentric axial dipole (‘bar magnet’). During normal polarity of the field the average magnetic north pole is at the geographic north pole, and a compass aligns along magnetic field lines. Historically, we refer to the north pole as the pole attracting the ‘north-seeking’ needle of a compass, but physically it is a south pole. During normal polarity, the inclination is positive (downward-directed) in the Northern Hemisphere and negative (upward-directed) in the Southern Hemisphere. Conversely, during reversed polarity, the compass needle points south, and the inclination is negative in the Northern and positive in the Southern Hemisphere.
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higher-order terms: the nondipole field. At any one time, the best-fitting geocentric dipole axis does not coincide exactly with the rotational axis of the Earth, but averaged over a few thousand years we may treat the dipole as both geocentric and axial. The most distinctive property of the Earth’s magnetic field is that it can reverse its polarity: the north and south poles interchange. Paleomagnetic studies of igneous rocks provided the first reliable information on reversals. In 1906, Brunhes observed lava flows magnetized in a direction approximately antiparallel to the present geomagnetic field, and suggested that this was caused by a reversal of the field itself, rather than by a self-reversal mechanism of the rock. In 1929, Matuyama demonstrated that young Quaternary lavas were magnetized in the same direction as the present field (normal polarity), whereas older lavas were magnetized in the opposite direction. His study must be regarded as the first magnetostratigraphic investgation. Initially, it was believed that the field reversed periodically, but as more (K/Ar dating plus paleomagnetic) results of lava flows became available, it became clear that geomagnetic reversals occur randomly. It is this random character that fortuitously provides the distinctive ‘fingerprints’ and gives measured polarity sequences their correlative value. A polarity reversal typically takes several thousands of years, which on geological timescales is short and can be taken as globally synchronous. The field itself is sign invariant: the same configuration of the geodynamo can produce either a normal or reversed polarity. What causes the field to reverse is still the subject of debate, but recent hypotheses suggests that lateral changes in heat flow at the core–mantle boundary play an important role. Although polarity reversals occur at irregular times, over geological time spans the reversal frequency can change considerably. For instance, the polarity reversal frequency has increased from approximately 1 My 1 some 80 My ago to 5 My 1 in more recent times. During part of the Creataceous, no reversals occurred at all from 110 to 80 Ma: the field showed a stable normal polarity during some 30 My. Such long periods of stable polarity are called Superchrons, and the one in the Cretaceous is also recognized as the Cretaceous Normal Quiet Zone in oceanfloor magnetic anomalies. The ancient geomagnetic field can be reconstructed from its recording in rocks during the geological past. Almost every type of rock contains magnetic minerals, usually iron oxide/hydroxides or iron sulfides. During the formation of rocks, these magnetic minerals (or more accurately: their magnetic domains) statistically align with the then ambient field, and will subsequently be ‘locked in,’
preserving the direction of the field as natural remanent magnetization (NRM): the paleomagnetic signal. The type of NRM depends on the mechanism of recording the geomagnetic signal, and we distinguish three basic types: TRM, CRM, and DRM. TRM, or thermoremanent magnetization, is the magnetization acquired when a rock cools through the Curie temperature of its magnetic minerals. Curie temperatures of the most common magnetic minerals are typically in the range 350–7001C. Above this temperature, the magnetic domains align instantaneously with the ambient field. Upon cooling, they are locked: the magnetic minerals carry a remanence that usually is very stable over geological time spans. Any subsequent change of the direction of the ambient field cannot change this remanence. Typically, TRM is acquired in igneous rocks. CRM, or chemical remanent magnetization, is the magnetization acquired when a magnetic mineral grows through a critical ‘blocking diameter’ or grain size. Below this critical grain size, the magnetic domains can still align with the ambient field; above it, the field will be locked and the acquired remanence may again be stable over billions of years. CRM may be acquired under widely different circumstances, e.g., during slow cooling of intrusive rocks, during metamorphosis or (hydrothermal) fluid migration, but also during late diagenetic processes such as weathering processes through formation of new magnetic minerals. A particularly important mechanism of CRM acquisition occurs in marine sediments during early diagenesis: depending on the redox conditions, iron-bearing minerals may dissolve, and iron may become mobile. If the mobilized iron subsequently encounters oxic conditions, it may precipitate again and form new magnetic minerals that then acquire a CRM. DRM, or detrital remanent magnetization, is the magnetization acquired when magnetic grains of detrital origin already carrying TRM or CRM are deposited. The grains statistically align with the ambient field as long as they are in the water column or in the soft water-saturated topmost layer of the sediment (Figure 2). Upon compaction and dewatering, the grains are mechanically locked — somewhere in a ‘lock-in depth zone’ — and will preserve the direction of the ambient field. The Magnetic Signal in Marine Sediments
In sediments, it is often assumed that the natural remanent magnetization is due to DRM. In the past, there has been some debate on the mechanisms of DRM acquisition, and the term pDRM or postdepositional DRM has often been used, because the
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sedimentary environment. Depending on the role of organic matter, existing magnetic minerals may dissolve and iron may be mobilized and precipitate elsewhere. These geochemical processes may results in the (partial) removal of the original paleomagnetic signal B and in the acquisition of CRM in a particular zone (much) later than the deposition of the sediment in this zone. Clearly, such processes may severely damage the Still water fidelity of the paleomagnetic record, and may offset the position of a reversal boundary by a distance in sediment that can correspond to a time of up to tens of thousands of years. However, in ‘suitable’ sediments Sediment /water interface the damage is usually restricted, and the paleomagnetic signal of such sediments may be considered as Bioturbation reliable and near-depositional. Suitable environments DRM lock-in depth are generally those with a sufficiently high sedimentation rate, a significant detrital input, and a preCompaction dominantly oxic environment. Therefore, it is often Oxic sediment O O O O Oxygen diffusion necessary to check the origin of the NRM using rock magnetic and geochemical methods. Precipitation of iron oxides As a rule, the total NRM is composed of different and CRM formation Anoxic sediment components. Ideally, the primary NRM — that oriDissolution of iron oxides ginating from near the time of deposition — has been conserved, but often this original signal is contaminated with ‘viscous’ remanence components, Fe Fe Fe Fe Iron diffusion (Sub)oxic sediment referred to as VRM. Such a VRM may result from partial realignment of ‘soft’ magnetic domains in the Figure 2 Depositional remanent magnetization (DRM) present-day field or from low-temperature oxidation acquired in sediments involves a continuum of physical and chemical processes. Detrital magnetic minerals (black) will be of magnetic minerals. It is generally easily removed statistically aligned along the ambient geomagnetic field (B) in through magnetic ‘cleaning,’ which consists of a still water and in the soft and bioturbated water-saturated routine laboratory treatment called demagnetization; sediment just below the sediment–water interface. Upon details can be found in any standard texbook on compaction and dewatering, the grains are mechanically paleomagnetism. Despite these pitfalls in sedimentlocked, preserving the direction of the field. Early diagenetic processes such as sulfate reduction may dissolve iron-bearing ary paleomagnetic records, sediments — in contrast minerals. Upon encountering a more oxic environment, iron may to igneous rocks — have in principle the distinctive precipitate as iron oxides, which will acquire a chemical remanent advantage of providing a continuous record of the magnetisation (CRM). The thus-acquired CRM in this layer may geomagnetic field, including the history of geohave a much later age than the depositional age of this layer. magnetic polarity reversals. Paleomagnetic studies of the sediments will then reveal the pattern of geolock-in zone has a certain depth. Sediment at this magnetic reversals during deposition. depth is slightly older than that the sediment–water interface. Hence, the acquisition of NRM is always slightly delayed with respect to sediment age. For The Geomagnetic Polarity Timescale practical purposes (in magnetostratigraphy) this usually has no serious consequences. Nevertheless, the (GPTS) concept of a purely detrital remanent magnetization an Surveys over the ocean basins carried out from the oversimplification of the real world. We therefore 1950s onward found linear magnetic anomalies, prefer to use the term depositional remanent magnet- parallel to mid-ocean ridges, using magnetometers ization (DRM) to refer to a continuum of physical and towed behind research vessels. During the early chemical processes that occur during and shortly after 1960s, it was suggested, and soon confirmed, that deposition. The acquisition of a DRM thus depends these anomalies resulted from the remanent magboth on detrital input of magnetic minerals and the netization of the oceanic crust. This remanence is new formation of such minerals in the sediment acquired during the process of seafloor spreading, through early diagenetic processes (Figure 2). Early when uprising magma beneath the axis of the middiagenesis is widespread and occurs in virtually every ocean ridges cools through its Curie temperature Turbulent water
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(Figure 3) in the ambient geomagnetic field, thus acquiring its direction and polarity. The continuous process of rising and cooling of magma at the ridge results in magnetized crust of alternating normal and reversed polarity that produces a slight increase or decrease of the measured field — the marine magnetic anomalies. It was also found that the magnetic anomaly pattern is generally symmetric on both sides of the ridge, and, most importantly, that it provides a wonderfully continuous ‘tape recording’ of the geomagnetic reversal sequence. A major step in constructing a time series of polarity reversals was taken in 1968 by Heirztler and co-workers. They used a long profile from the southern Atlantic Ocean and, extrapolating from a known age for the lower boundary of the Gauss Chron (Figure 4), they constructed a geomagnetic polarity timescale under the assumption of constant 2 Ma
1 Ma
spreading. Subsequent revisions mostly used the anomaly profile of Heirtzler, often adding more detail from other ocean basins, and appending additional calibration points. These calibration points are derived from sections on land, which, first, must contain a clear fingerprint of magnetic reversals that can be correlated to the anomaly profile, and second, contain rocks that can be reliably dated by means of radiometric methods. The GPTS is then derived by linear interpolation of the anomaly pattern between these calibration points, again under the assumption of constant spreading rate between those points. The development of the GPTS (Figure 4) shows increasing detail and gradually improved age control. Periods of a predominant (normal or reversed) polarity are called chrons, and the four youngest ones are named after individuals: Brunhes (normal), who suggested field reversal; Matuyama (reversed),
Present
1 Ma
2 Ma
Model profile
Observed profile
Mid oceanic ridge
Sediment
Uprising magma and ‘locking in’ of magnetic polarity
Crust Lithosphere
Figure 3 Formation of marine magnetic anomalies during seafloor spreading. The oceanic crust is formed at the ridge crest, and while spreading away from the ridge it is covered by an increasing thickness of oceanic sediments. The black (white) blocks of oceanic crust represent the original normal (reversed) polarity thermoremanent magnetization (TRM) acquired upon cooling at the ridge. The black and white blocks in the drill holes represent normal and reversed polarity DRM acquired during deposition of the marine sediments. The model profile (grey) represents computed magnetic anomalies produced by the block model of TRM polarity (top); the observed profile (dark) is the observed sea-level magnetic anomaly profile due to the magnetized oceanic crust.
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APTS, 2000
Cande and Kent, 1992
Berggren et al., 1985
LaBreque et al., 1977
Heirtzler et al., 1968
Cox et al., 1963
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Brunhes
0
Matuyama
1
Gauss
2
3
Gilbert
Age (Ma)
4
5
6
7
8 Figure 4 Development of the geomagnetic polarity timescale (GPTS) through time. The initial assumption of periodic behavior (in 1963) was soon abandoned as new data became available. The first modern GPTS based on marine magnetic anomaly patterns was established in 1968 by Heirtzler and co-workers. Subsequent revisions show improved age control and increased resolution. A major breakthrough came with the astronomical polarity timescale (APTS), in which each individual reversal is accurately dated.
who proved this; Gauss (normal), who mathematically described the field; and Gilbert (reversed), who discovered that the Earth itself is a huge magnet. Chrons may contain short intervals of opposite
polarity called subchrons, which are named after the locality where they were discovered; for example, the normal Olduvai subchron within the Matuyama reversed chron is named after Olduvai Gorge in
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Tanzania, and the Kaena reversed subchron within the Gauss normal chron after Kaena Point on Hawaii. Older chrons were not named but numbered, according to the anomaly numbers earlier given by Heirtzler, which has led to a confusing nomenclature of chrons and subchrons. A major step forward was taken by Cande and Kent in 1992, who thoroughly revised the magnetic anomaly template over the last 110 My. They constructed a synthetic flow line in the South Atlantic, using a set of chosen anomalies that were taken as tie points. The intervals between these tie points were designated as category I intervals. On these intervals they projected (stacks of) the bestquality profiles surveyed in this ocean basin, providing category II intervals. Since the spreading of the Atlantic is slow, they subsequently filled in the category II intervals with high-resolution profiles from fast-spreading ridges (their category III). This enabled them to include much more detail on short polarity intervals (or subchrons); see, for instance, the increase of detail around 7 Ma in Figure 4. Very short and low-intensity anomalies still have an uncertain origin. They may represent very short subchrons of the field, as has been proven for some of them (e.g., the Cobb Mt. subchron at 1.21 Ma, or the Re´union subchron at 2.13–2.15 Ma), or may just represent intensity fluctuations of the geomagnetic field causing the oceanic crust to be less (or more) strongly magnetized. Because of their uncertain or unverified nature, these were called cryptochrons. In addition, Cande and Kent developed a consistent (sub)chron nomenclature that is now used as the standard. In total, they used nine calibration points, but they made a break with tradition by using, for the first time, an astronomically dated age tie point for the youngest one: the Gauss/Matuyama boundary. The correlation of the GPTS to global biostratigraphic zonations is covered extensively by Berggren et al. (1995). The template of magnetic anomaly patterns from the ocean floor has remained central for constructing the GPTS from the late Cretaceous onward (110–0 Ma). Only recently, the younger part of the GPTS has been based on direct dating of each individual reversal through the use of orbitally tuned timescales. In their most recent version of the GPTS, Cande and Kent included the astronomical ages for all reversal boundaries for the past 5.3 My. The CK95 or Cande and Kent (1995) geomagnetic polarity timescale is at present the most widely used standard. Polarity timescales for the Mesozoic rely on some of the oldest magnetic anomaly profiles, down to the late Jurassic, and on dated magnetostratigraphies of sections on land.
The Astronomical Polarity Timescale (APTS) The latest development in constructing a GPTS comes from orbital tuning of the sediment record; for details see the article Orbitally Tuned Timescales. It differs essentially from the conventional GPTS in the sense that each reversal boundary — or any other geological boundary for that matter, e.g. biostratigraphic datum levels or stage and epoch boundaries — is dated individually. This has provided a breakthrough in dating of the geological record and has the inherent promise of increasing understanding of the climate system, since cyclostratigraphy and subsequent orbital tuning rely on decphering and understanding environmental changes driven by climate change, which in turn is orbitally forced. The fact that the age of each reversal is directly determined, rather than interpolated between calibration points, has important consequences for (changes in) spreading rates of plat pairs. Rather than having to assume constant spreading rates between calibration points, one can now accurately determine these rates, and small changes therein. Indeed, Wilson found that the use of astronomical ages resulted in very small and physically realistic spreading rate variations. As a result, the discrepancy in between plate motion rates from the global plate tectonic model (NUVEL-1) and those derived from geodesy has become much smaller. Meanwhile NUVEL-1 has been updated (to NUVEL-1A) to incorporate the new astronomical ages. Another application is the dating of Pleistocene, Pliocene, and Miocene, and older stage boundaries, many of which have been defined in the Mediterranean. The availability of a good astrochronology has effectively become a condition for the definition of a Global Boundary Stratotype Section and Point (GSSP). An example is shown in Figure 5, showing the Tortonian–Messinian GSSP that has recently been defined in the Atlantic margin basin of western Morocco. Perhaps one of the most promising areas of the application of astrochronology is in the bed-to-bed correlation of the two different realms of oceans and continent. Climate forcing may be expected to have a different expression in the different realms because of the different nature of their sedimentary environments. A recently established and refined orbital timescale for the loess sequences of northern China relies upon the correlation of detailed monsoon records to the astronomical solutions and the oceanic oxygen isotope records. An important finding was that the straightforward use of magnetostratigraphy and correlation to the GPTS cannot provide a sufficiently accurate age model for comparison with the
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GEOMAGNETIC POLARITY TIMESCALE
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Figure 5 Magnetostratigraphy, cyclostratigraphy, and astrochronology of the marine Oued Akrech section from the Atlantic margin of Morocco, which defines the Tortonian–Messinian Global Stratotype Section and Point (GSSP). The sedimentary cycles represent a cyclically changing environment that correlates with variations in insolation. Insolation is strongly related to climatic precession (upper left panel), which induces cyclic changes in seasonal contrast, reflected one-on-one in the sedimentary cycles. (After Hilgen FJ et al. (2000) Episodes 23(3): 172–178.)
ocean record, since the analysis of the astrochronological framework demonstrates considerable downward displacement of reversal boundaries because of delayed lock-in of the NRM. With the new chronology and its direct correlation to the oceanic record, it is now possible to analyse terrestrial paleomonsoon behavior for the past 2.6 My and compare it to climate proxies from the marine realm. This may give important information, for instance, on leads and lags of various systems in response to climate change, on phase relations with insolation, or on the relation between global ice volume and monsoonal climate.
See also Aeolian Inputs. Magnetics. Monsoons, History of. Paleoceanography, Climate Models in. Paleoceanography: Orbitally Tuned Timescales.
Further Reading Berggren WA, Kent DV, Aubry MP, and Hardenbol J (1995) Geochronology, Time Scales and Global Stratigraphic Correlations: A Unified Temporal Framework for an Historical Geology. Society of Economic Paleontologists and Mineralogist, Special Volume 54.
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Butler RF (1992) Paleomagnetism, Magnetic Domains to Geologic Terranes. Boston: Blackwell Scientific. Cande SC and Kent DV (1992) A new geomagnetic polarity Time-Scale for the Late Cretaceous and Cenozoic. Journal of Geophysical Research 97: 13917--13951. Cande SC and Kent DV (1995) Revised calibration of the Geomagnetic Polarity Time Scale for the Late Cretaceous and Cenozoic. Journal of Geophysical Research 100: 6093--6095. Gradstein FM, Agterberg FP, Ogg JG, et al. (1995) A Triassic, Jurassic and Cretaceous time scale. In: Berggren WA, Kent DV, Aubry M and Hardenbol J (eds) Geochronology, Time Scales and Stratigraphic Correlation. Society of Economic Paleontologists and Mineralogist, Special Volume 54: 95–128. Hilgen FJ, Krijgsman W, Langereis CG, and Lourens LJ (1997) Breakthrough made in dating of the geological record. EOS, Transactions of the AGU 78: 285--288. Heirtzler JR, Dickson GO, Herron EM, Pitman WC III, and Le Pichon X (1968) Marine magnetic anomalies,
geomagnetic field reversals, and motions of the ocean floor and continents. Journal of Geophysical Research 73: 2119--2136. Heslop D, Langereis CG, and Dekkers MJ (2000) A new astronomical time scale for the loess deposits of Northern China. Earth and Planetory Science Letters 184: 125--139. Maher BA and Thompson R (1999) Quaternary Climates, Environments and Magnetism. Cambridge: Cambridge University Press. Opdyke ND and Channell JET (1996) Magnetic Stratigraphy. San Diego: Academic Press. Shackleton NJ, Crowhurst SJ, Weedon GP, and Laskar J (1999) Astronomical calibration of Oligocene–Miocene time. Philosophical Transactions of the Royal Society of London, A 357: 1907--1929. Wilson DS (1993) Confirmation of the astronomical calibration of the magnetic polarity time scale from seafloor spreading rates. Nature 364: 788--790.
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GEOMORPHOLOGY C. Woodroffe, University of Wollongong, Wollongong, NSW, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1142–1148, & 2001, Elsevier Ltd.
Introduction Geomorphology is the study of the form of the earth. Coastal geomorphologists study the way that the coastal zone, one of the most dynamic and changeable parts of the earth, evolves, including its profile, plan-form, and the architecture of foreshore, backshore, and nearshore rock and sediment bodies. To understand these it is necessary to examine wave processes and current action, but it may also involve drainage basins that feed to the coast, and the shallow continental shelves which modify oceanographic processes before they impinge upon the shore. Morphodynamics, study of the mutual co-adjustment of form and process, leads to development of conceptual, physical, mathematical, and simulation models, which may help explain the changes that are experienced on the coast. In order to understand coastal variability from place to place there are a number of boundary conditions which need to be considered, including geophysical and geological factors, oceanographic factors, and climatic constraints. At the broadest level plate-tectonic setting is important. Coasts on a plate margin where oceanic plate is subducted under continental crust, such as along the western coast of the Americas, are known as collision coasts. These are typically rocky coasts, parallel to the structural grain, characterized by seismic and volcanic activity and are likely to be uplifting. They contrast with trailing-edge coasts where the continental margin sits mid-plate, which are the locus of large sedimentary basins. Smaller basins are typical of marginal sea coasts, behind a tectonically active island arc. The nature of the material forming the coast is partly a reflection of these broad plate-tectonic factors. Whether the shoreline is rock or unconsolidated sediment is clearly important. Resistant igneous or metamorphic rocks are more likely to give rise to rocky coasts than are those areas composed of broad sedimentary sequences of clays or mudstones. Within sedimentary coasts sandstones may be more resistant than mudstones. Rocky coasts tend to be relatively
resistant to change, whereas coasts composed of sandy sediments are relatively easily reshaped, and muddy coasts, in low-energy environments, accrete slowly. The relative position of the sea with respect to the land has changed, and represents an important boundary condition. The sea may have flooded areas which were previously land (submergence), or formerly submarine areas may now be dry (emergence). The form of the coast may be inherited from previously subaerial landforms. Particularly distinctive are landscapes that have been shaped by glacial processes, thus fiords are glacially eroded valleys, and fields of drumlins deposited by glacial processes form a prominent feature on paraglacial coasts. The oceanographic factors that shape coasts include waves, tides, and currents. Waves occur as a result of wind transferring energy to the ocean surface. Waves vary in size depending on the strength of the wind, the duration for which it has blown and the fetch over which it acts. Wave trains move out of the area of formation and are then known as swell. The swell and wave energy received at a coast may be a complex assemblage generated by specific storms from several areas of origin. Tides represent a largewavelength wave formed as a result of gravitational attractions of the sun and moon. Tides occur as a diurnal or semidiurnal fluctuation of the sea surface that may translate into significant tidal currents particularly in narrow straits and estuaries. In addition storm-generated surges and tsunamis may cause elevated water levels with significant geomorphological consequences. The coast is shaped by these oceanographic processes working on the rocks and unconsolidated sediments of the shoreline. Climate is significant in terms of several factors. First, wind conditions lead to generation of waves and swell, and may blow sand into dunes along the backshore. Climate also influences the rate at which weathering and catchment processes operate. In addition, regional-scale climate factors such as monsoonal wind systems and the El Nin˜o Southern Oscillation phenomenon demonstrate oscillatory behavior and may reshape the coast seasonally or interannually. Gradual climate change may mean that shoreline fluctuations do not revolve around stationary boundary conditions but may exhibit gradual change themselves. In this respect the impact of perceived global climate change during recent decades as a result of human-modified environmental factors is likely to be
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GEOMORPHOLOGY
Sea level
Sea level
(B)
(A)
Figure 1 Schematic representation of the original woodcut illustrations by Charles Darwin showing (A) the manner in which he deduced that fringing reefs around a volcanic island would develop into barrier reefs (dotted), and (B) barrier reefs would develop into an atoll, as a result of subsidence and vertical reef growth.
felt in the coastal zone. Of particular concern is anticipated sea-level rise which may have a range of geomorphological effects, as well as other significant socioeconomic impacts.
History Geomorphology has its origins in the nineteenth century with the results of exploration, and the realization that the surface of the earth had been shaped over a long time through the operation of processes that are largely in operation today (uniformitarianism). The observations by Charles Darwin during the voyage of the Beagle extended this view, particularly his remarkable deduction that fringing reefs might become barrier reefs which in turn might form atolls as a result of gradual subsidence of volcanic islands combined with vertical reef growth (Figure 1). In the first part of the twentieth century geomorphology was dominated by the ‘geographical cycle’ of erosion of William Morris Davis who anticipated landscape
denudation through a series of stages culminating in peneplanation, and subsequent rejuvenation by uplift. This highly conceptual model, across landscapes in geological time, was also applied to the coast by Davis who envisaged progressive erosion of the coast reducing shoreline irregularities with time (Figure 2). Such landscape-scale studies were extended by Douglas Johnson who emphasized the role of submergence or emergence as a result of sea-level change (Figure 3). The Davisian view was reassessed in the second half of the twentieth century with a greater emphasis on process geomorphology whereby studies focused on attempting to measure rates of process operation and morphological responses to those processes, reflecting ideas of earlier researchers such as G. K. Gilbert. The concept of the landscape as a system was examined, in which coastal landforms adjusted to equilibrium, perhaps a dynamic equilibrium, in relation to processes at work on them. Studies of
i ii iii (A) Initial
Figure 2 Schematic representation of the planform stages that W. M. Davis conceptualized through which a shoreline would progress from an initial rugged form (1) through maturity (2) to a regularized shoreline (3) that is cliffed (hatched) and infilled with sand (stippled). He envisaged this in parallel to the geographical cycle of erosion by which mountains (like the initial shoreline form, 1) would be reduced to a peneplain (like the solid ultimate shoreline, 3).
(B) Modern
Figure 3 Schematic representation of (A) the initial shoreline form envisaged by Douglas Johnson for an area in Marthas Vineyard, New England, and (B) the modern regularized, form of the shore.
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GEOMORPHOLOGY
sediment movement and the adjustment of beach shape under different wave conditions were typical.
Scales of Study Coastal geomorphology now studies landforms and the processes that operate on them at a range of spatial and temporal scales (Figure 4). At the smallest scale geomorphology is concerned with an ‘instantaneous’ timescale where the principles of fluid dynamics apply. It should be possible to determine details of sediment entrainment, complexities of turbulent flow and processes leading to deposition of individual bedform laminae. The laws of physics apply at these scales, though they may operate stochastically. In theory, behavior of an entire embayment could be understood; in practice studies simply cannot be undertaken at that level of detail and broad extrapolations based on important empirical relationships are made. The next level of study is the ‘event’ timescale, which may cover a single event such as an individual storm or an aggregation of several lesser events over a year or more. The mechanistic relationships from instantaneous time are scaled up in a deterministic or empirical way to understand the operation of coasts at larger spatial and temporal scales. Thus, stripping of a beach during a single storm, and the more gradual reconstruction to its original state under ensuing calmer periods can be observed. Time taken for reaction to the event, and relaxation back to a more ambient state may be known from surveys of beaches, enabling definition of a ‘sweep zone’ within which the beach is regularly active.
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At a larger scale of operation, the coastal geomorphologist is interested in the way in which coasts change over timescales that are relevant to societies and at which coastal-zone managers need to plan. This is the ‘engineering’ timescale, involving several decades. It is perhaps the most difficult timescale on which to understand coastal geomorphology and to anticipate behavior of the shoreline. The largest timescales are geological timescales. Studies over geological timescales are primarily discursive, conceptual models which recognize that boundary conditions, including climate and the rate of operation of oceanographic processes, change. It is clear that sea level itself has changed dramatically over millennia as a result of expansion and contraction of ice sheets during the Quaternary ice ages, and so the position of the coastline has changed substantially between glaciations and interglaciations.
Models of Coastal Evolution
Space
There has been considerable improvement in understanding the long-term development of coasts as a result of significant advances in paleoenvironmental reconstruction and geochronological techniques (especially radiometric dating). Incomplete records of past coastal conditions may be preserved, either as erosional morphology (notches, marine terraces, etc.) or within sedimentary sequences. Although this record is selective, reconstructions of Quaternary paleoenvironments, together with interpretation of geological sequences in older rocks, have enabled the formulation of geomorphological models based on sedimentary evidence. In the case of deltas, where there may be important hydrocarbon reserves, the complementary development of geological models GEOLOGICAL Global and study of modern deltas has led to better underReef standing of process and response at longer timescales structure ENGINEERING than those for which observations exist. Cliff Unconsolidated sediments are the key to coastal Regional retreat morphodynamics because coasts change through Delta EVENT dynamics erosion, transport, and deposition of sediment. Study of coastal systems has led to many insights, parReef-flat morphology ticularly in terms of the various pathways through Mudflat Local sedimentation which sediment may move within a coastal comBeach states partment or circulation cell. However, it is clear that INSTANTANEOUS a completely reductionist approach to coastal geomm Seconds Years Millennia morphology will not lead to understanding of all Time components and the way in which they interact. Empirical relationships and the presumed determinFigure 4 Representation of space and timescales appropriate istic nature of sediment response to forcing factors for the study of coastal geomorphology and schematic representation of some of the examples discussed in the context of remain incomplete and are all too often formulated instantaneous, event, engineering, and geological scales of on the basis of presumed uniform sediment sizes enquiry. or absence of biotic influence and are ultimately
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GEOMORPHOLOGY
Bore Reflected standing wave
Breaking wave
(B)
geomorphology of each of these behaves differently. Processes operate at different rates, transitions between different states occur over different timescales and the significance of antecedent conditions varies. Rocky Coasts
Bore (A)
Breaking wave
(C)
Figure 5 The morphology of cliffs and shore platforms. Plunging cliffs (A) reflect wave energy and are generally not eroded. Shore platforms dissipate wave energy and may be either (B) ramped, or (C) subhorizontal. Change in these systems tends to occur at geological (hard rock) or engineering (soft rock) timescales.
unrealistic. More recently coastal systems have been investigated as nonlinear dynamical systems, because relationships are not linear and behavior is potentially chaotic. Nonlinear dynamical systems behave in a way that is in part dependent upon antecedent conditions. This is well illustrated by studies of beach state; beach and nearshore sediments are particularly easily reshaped by wave processes, thus a series of characteristics typical of distinct beach states can be recognized (see Figure 6). However, the beach is only partly a response to incident wave conditions, its shape being also dependent on the previous shape of the beach which was in the process of adjusting to wave conditions incident at that time. Coastal systems are inherently unpredictable. However, they may operate within a broad range of conditions with certain states being recurrent. Chaotic systems may tend towards self-organization, for instance patterns of beach cusps characteristic along low-energy beaches which reflect much of the wave energy may adopt a self-organized cuspate morphology in which swash processes, sediment sorting, and form are balanced. Models may be developed based on patterns of change inferred over geological timescales but consistent with mechanistic processes known to operate over lesser event timescales. Simulation modelling is not intended to reconstruct coastal evolution exactly, but becomes a tool for experimentation and extrapolation within which broad scenarios of change can be modeled and sensitivity to parameterization of variables examined. Coasts may be divided into rocky coasts, sandy coasts, and muddy coasts, and coral reefs and deltas and estuaries can be differentiated. The
Rocky coasts are characterized by sea cliffs, especially on tectonically active, plate-margin coasts where there are resistant rocks. Cliffed coastlines in resistant rock appear to adopt one of two forms (states): plunging cliffs where a vertical cliff extends below sea level, and shore platforms where a broad bench occurs at sea level in front of a cliff (Figure 5). Erosion of cliffs occurs where the erosional force of waves exceeds the resistance of the rocks, and where sufficient time has elapsed. Plunging cliffs occur where the rock is too resistant to be eroded. The vertical face results in a standing wave which reflects wave energy, so that there is little force to erode the cliff at water level. Waves exert a greater force if they break, or if they are already broken. This can only occur if the water depth is shallow offshore from the cliff face in which case the increased energy from the breaking waves is also able to entrain sediment from the floor (or rock fragments quarried from the foot of the cliff). This process of erosion accelerates through a positive feedback cutting a shore platform. Shore platforms thus develop in those situations where the erosive force of waves exceeds resistance of the rock. A platform widens as a result of erosion at the foot of the cliff behind the platform. The cliff oversteepens, leading to toppling and fall of detritus onto the rear of the platform, which slows further erosion of the cliff face until that talus has been removed. A series of such negative feedbacks slow the rate at which platforms widen over time, and there is often considerable uniformity of platforms up to a maximum width in any particular lithological setting. Shore platforms adopt either a gradually sloping ramped form, or a subhorizontal form often with a seaward rampart (Figure 5). There has been much discussion as to the relative roles of wave and subaerial processes in the formation of these platforms. Platforms in relatively sheltered locations appear to owe their origin to processes of water-layer leveling (physiochemical processes in pools which persist on platforms at low tide) or wetting and drying and its weakening of the rock. In other cases wave quarrying and abrasion are involved. In many cases both processes may be important. Other platforms may be polygenetic with inheritance from former stands of sea level (reflecting antecedent conditions).
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GEOMORPHOLOGY
(A) Dissipative
(B) Intermediate
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(C) Reflective
Plunging waves
Spilling waves
Accreted beach
Surging waves
Nearshore bar (D)
Storm
Beach volume
C
A Time
Figure 6 Beach state may be (A) dissipative, where high wave-energy impinges on low-gradient beaches; (B) intermediate (in which several states are possible); or (C) reflective, where low wave-energy surges onto steep or coarse beaches. Change between beach states occurs at event timescales. (D) A beach may react rapidly to a perturbation such as a storm changing from reflective state (C) to more dissipative state (A) but then change much more slowly through intermediate states, readjusting towards state (C) again unless further disrupted.
Coral Reefs
Sandy Coasts
A group of coastlines that are of particular interest to the geomorphologist are those formed by coral reefs. Corals are colonial animals that secrete a limestone exoskeleton that may form the matrix of a reef. Coral reefs flourish in tropical seas in high-energy settings where a significant swell reaches the shoreline. In these circumstances the reef attenuates much of the wave energy such that a relatively quiet-water environment occurs sheltered behind the reef. Not only is this reef ecosystem of enormous geomorphological significance in terms of the solid reef structure that it forms, but in addition, the carbonate reef material breaks down into calcareous gravels, sands and muds which form the sediments that further modify these coastlines. On the one hand reefs can be divided into fringing reefs, barrier reefs, and atolls, distinct morphological states that form part of an evolutionary sequence as a result of gradual subsidence of the volcanic basement upon which the reef established (see Figure 1). This powerful deduction by Darwin relating to reef structure and operating over geological timescales, has been generally supported by drilling and geochronological studies on mid-plate islands in reef-forming seas. On the other hand, the surface morphology of reefs is extremely dynamic with rapid production of skeletal sediments and their redistribution over event timescales, with the landforms of reefs responding to minor sea-level oscillations, storms, El Nin˜o, coral bleaching, and other perturbations.
Beaches form where sandy (or gravel) material is available forming a sediment wedge at the shoreline. The beach is shaped by incident wave energy, and can undergo modification particularly by formation and migration of nearshore bars which in turn modify the wave-energy spectrum. Various beach states can be recognized across a continuum from beaches which predominantly reflect wave energy, and those which dissipate wave energy across the nearshore zone. Reflective beaches are steep, waves surge up the beach and much of the energy is reflected, and the beach face may develop cusps. Dissipative beaches are much flatter, and waves spill before reaching the shoreline (Figure 6). During a storm, sand is generally eroded from the beach face and deposited in the nearshore, often forming shore-parallel or transverse bars. Waves consequently break on the bars and energy is lost, the form modifying the process in a mutual way. It may be possible to recognize a series of beach states intermediate between reflective and dissipative and any one beach may adopt one or several beach states over time (Figure 6). Beach state is clearly modified by incident wave conditions, but the rate of adjustment between states takes time, and a beach is also partly a function of antecedent beach states. Although erosion and redeposition of sand is the way in which the nearshore adjusts, there may also be long-term storage of sediment. In particular broad, flat beaches may develop dunes behind them.
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In other cases a sequence of beach ridges may develop. Geomorphological changes in state over geological timescales represented by beach-ridge plains may relate to variations in supply of sediment to the system (perhaps by rivers) as well as variations in the processes operating. Deltas and Estuaries
Where rivers bring sediment to the coast, deltas and estuaries can develop. In this case the sediment budget of the shoreline compartment or the receiving basin is augmented and there is generally a positive sediment budget. Deltas are characterized by broad wedges of sediment deposition. Delta morphology tends to reflect the processes that are dominant (Figure 7). Where wave energy is low and tides are minor, it is primarily river flows which account for sedimentation patterns. River flow may be likened to a jet, influenced by inertial forces, friction with the basin floor or buoyancy where there are density differences between outflow and the water of the receiving basin. Wave action tends to smooth the shoreline and a wave-dominated deltaic shoreline will be characterized by shore-parallel bars or ridges. Where riversupplied sediment is relatively low in volume, wave action may form a sandy barrier along the coast and rivers may supply sediment only to lagoons formed behind these barriers. Such is the case for much of the barrier-island shoreline of the eastern shores of the Americas; where barrier islands may be continually reworked landwards during relative sea-level rise. On the other hand if sea level is stable, a stable sand barrier is likely to form closing each
River
Wave
Tide
Figure 7 Deltas adopt a variety of forms, but morphology appears to reflect the relative balance of river, wave, and tide action. The broad morphology of the delta tends to be digitate where river-dominated, shore-parallel where wave-dominated, and tapering where tide-dominated.
embayment or creating a barrier estuary as in the case of coastal lagoons in Asia and Australasia. Estuaries are embayments which are likely to infill incrementally either through the deposition of riverborne sediments or through the influx of sediment from seaward by wave or tidal processes (Figure 8). The sediment from seaward may either be derived from the shelf, or from shoreline erosion. Tidal processes differ from river processes; they are bi-directional, flowing in during flood tide and out during ebb, and the flow is forced by the rising level of the sea. Small embayments are flooded and drained by a tidal prism (the volume of water between low and high tide). Longer estuaries, on the other hand, may have a series of tidal waves which progress up them. Where the tidal range is large this may flow as a tidal bore. Tide-dominated estuarine channels adopt a distinctive tapered form, with width (and depth) decreasing from the mouth upstream (Figure 8). Wave and tidal processes tend to shape deltas and estuaries to varying degrees depending on their relative operation. Many of the deltaic-estuarine processes operate over cycles of change. The hydraulic efficiency of distributaries decreases as they lengthen, until an alternative, shorter course with steeper hydraulic gradient is adopted. The abandoned distributaries of deltas often become tidally dominated with sinuous, tapering tidal creeks dominating what may be a gradually subsiding abandoned delta plain. Wave processes, in wavedominated settings, smooth and rework the abandoned delta shore, often forming barrier islands. Muddy Coasts
Muddy coasts are associated with the lowest energy environments. Mud banks may occur in high-energy, wave-exposed settings where large volumes of mud are supplied to the mouth of large rivers. Thus longshore drift north west of the Amazon, and around Bohai Bay downdrift from the Yellow River, enables mud-shoal deposition in open-water settings. Elsewhere mud flats are typical of sheltered settings, within delta interdistributary bays, around coastal lagoons, etc. These muddy environments are likely to be colonized by halophytic vegetation. Mangrove forests occur in tropical settings, whereas salt marsh occurs in higher latitudes, often extending into tropical areas also. These coastal wetlands further promote retention of fine-grained sediment. The muddy environments are areas of complex hydrodynamics and sedimentation. Sedimentation is likely to occur with a negative feedback such that as the tidal wetlands
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GEOMORPHOLOGY
(A) Wave-dominated Tidal delta
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(B) Tide-dominated
Wetland
Wetland
Fluvial delta
Halopytes River Send shoal River
Sand barrier
Saline mudflats
Former shoreline Mud basin
Figure 8 Estuaries are broad embayments which may adopt a wide range of morphologies. They tend to show sand-barrier accumulation where wave-dominated (A), but be prominently tapering where tide-dominated (B). Estuaries are generally sediment sinks with the rates and patterns of infill reflecting the relative dominance of river, wave, and tide processes and sediment sources.
accrete sediment and as the substrate is elevated, they are flooded less frequently and therefore sedimentation decelerates. Boundary conditions, particularly sea level, are likely to vary at rates similar to the rate of sedimentation and prograded coastal plains contain complex sedimentary records of changes in ecological and geomorphological state.
See also
Conclusion
Boyd R, Dalrymple R, and Zaitlin BA (1992) Classification of clastic coastal depositional environments. Sedimentary Geology 80: 139--150. Carter RWG and Woodroffe CD (eds.) (1994) Coastal Evolution: Late Quaternary Shoreline Morphodynamics. Cambridge: Cambridge University Press. Cowell PJ and Thom BG (1994) Morphodynamics of coastal evolution. In: Carter RWG and Woodroffe CD (eds.) Coastal Evolution: Late Quaternary Shoreline Morphodynamics, pp. 33--86. Cambridge: Cambridge University Press. Darwin C (1842) The Structure and Distribution of Coral Reefs. London: Smith Elder. Davis RA (1985) Coastal Sedimentary Environments. New York: Springer-Verlag. Johnson DW (1919) Shore Processes and Shoreline Development. New York: Prentice Hall. Trenhaile AS (1997) Coastal Dynamics and Landforms. Oxford: Clarendon Press. Wright LD and Thom BG (1977) Coastal depositional landforms: a morphodynamic approach. Progress in Physical Geography 1: 412--459.
Coastal geomorphology is the study of the evolving form of the shoreline in response to mutually adjusting processes acting upon it. It spans instantaneous and event timescales, over which beaches respond to wave energy, through engineering and geological timescales, over which deltas build seaward or switch distributaries, sea level fluctuates, and cliff morphology evolves. The morphology (state) of the coast changes in response to perturbations, particularly extreme events such as storms, but also thresholds within the system (as when a cliff oversteepens and falls, or a distributary lengthens and then switches). Human action may also represent a perturbation to the system. In each coastal setting, the influence of human modifications is being felt. Thus there are fewer coasts which are not in some way influenced by society. As anthropogenic modification of climate and sea level occurs at a global scale, the human factor increasingly needs to be given prominence in coastal geomorphology.
Beaches, Physical Processes Affecting. Coral Reefs. Rocky Shores. Salt Marshes and Mud Flats.
Further Reading
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GEOPHYSICAL HEAT FLOW C. A. Stein, University of Illinois at Chicago, Chicago, IL, USA R. P. Von Herzen, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1149–1157, & 2001, Elsevier Ltd.
Introduction Heat flow is a directly measurable parameter at Earth’s surface that provides important constraints on the thermal state of its interior. When combined with other data, it has proved useful for models of Earth’s origin, structure, composition, and convective motion of at least the upper mantle. Because oceanic crust is mostly thin (about 5–8 km) and deficient in heat-producing radioactive elements compared to continental crust, marine heat flow measurements are particularly illuminating for the cooling and dynamics of the oceanic lithosphere. Indeed, they have provided an important constraint in the development and refinement of the plate tectonics paradigm over the past three decades. Initially influenced by concepts of Sir Edward Bullard, the first successful marine heat flow measurements obtained shortly after the end of World War II were among the first applications of remote electronics in the deep sea. Other than a general understanding of seafloor topography, little information on the nature of the Earth beneath the oceans was then available, so theories to explain the measurements were largely unconstrained. Probably the most important initial finding was that the mean oceanic heat flux was not much different from the mean of the available continental values, contrary to conventional expectations of that era. However, the entire field of marine geophysics (especially seismic reflection and refraction, and geomagnetism) was also being transformed by instrumentation development, and the availability of ships as seagoing platforms gave rise to a ‘golden age’ of ocean exploration between about 1950 and 1970. The large-scale (>1000 km) variations of marine heat flux were shown to be consistent with the plate tectonic paradigm that explained the mean depth and age of much of the ocean floor. Heat flow is now useful for studies of seafloor tectonics and other marine geophysical problems.
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Marine heat flow measurements now number more than 15 000, and with the exception of high latitude regions, are distributed over most of Earth’s main ocean basins (Figure 1). Values range over about two orders of magnitude (about 10–1000 mW m2), with a general systematic spatial distribution at large (41000 km) scales but also a large variability at small (o50 km) scales caused by pervasive hydrothermal circulation in mostly young (o50– 70 million years (My)) ocean crust. The spatial distribution and magnitudes of marine heat flow values may be used to deduce the geometry and intensity of such circulation.
Measurements and Techniques Instrumentation and Development
Techniques for marine measurements were developed with the realization that the temperature of the deep (42 km) ocean is mostly horizontally stratified, and relatively constant over long time scales (4100 years). In this situation, the heat flux from the Earth’s interior is reflected in a relatively uniform temperature gradient with depth beneath a flat seafloor. Heat flow is the product of the magnitude of this temperature gradient and the thermal conductivity of the material over which it is measured. Gradients are commonly measured using a vertically oriented probe with temperature sensors surmounted by a weight that is lowered from a ship to penetrate the relatively soft sediments that cover most of the seafloor. Thermal conductivity is either measured with in situ sensors during the seafloor penetration by transient heating experiments or aboard ship on sediment cores. For instrumental simplicity, the initial heat flow equipment used only two thermal sensors at either end of a 3–4 m long probe with in situ analog recording techniques on paper or film. Subsequently the number of sensors has been increased, due to a desire to measure non-linear temperature gradients and thermal conductivity that may vary with depth. The wide range of measured thermal gradients and the developments in solid-state electronics have made digital recording the present standard. ‘Pogo’ measurements, i.e., multiple penetrations on the same station, combined with nearly real-time acoustic telemetry of raw data have proven useful for investigation of small-scale (o10–20 km) marine
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Figure 1 Locations of marine heat flow measurements (dots). Thin lines show the location of the major plate boundaries.
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heat flow variability, the origins of which are still not fully understood. It was necessary for the ‘Bullard’ probe to remain undisturbed in the seafloor for most of one hour to approach thermal equilibrium with the sediments because of its relatively large diameter (about 2– 3 cm). This probe has been largely supplanted by either much smaller (about 3 mm diameter) individual probes mounted in outrigged fashion on a larger (‘Ewing’) probe to attain deep (5–6 m) penetration, or a ‘violin-bow (Lister)’ probe consisting of a small (about 1 cm) sensor string mounted parallel to, but separate from, the main strength member (Figure 2). Both are capable of in situ thermal conductivity measurements, utilizing either a constant heat source or a calibrated pulse (approximating a delta function) after the gradient measurement. The measurement time in the seafloor is 15–20 min. Depending mostly on the
desired spacing during pogo operations, the mean time between measurement is 1–2 h. Even with acoustic data telemetering, battery life can be 2–3 days, thereby allowing many measurements during a single lowering. Real-time ship location accuracy now approaches a few meters with differential Global Positioning System navigation, although the uncertainty of the probe location during pogo operations in normal ocean depths (4–5 km) is typically 200 m or more. Hence 1–2 km is a useful minimum spacing between pogo penetrations unless seafloor acoustic transponder navigation is employed for higher accuracy (about 10 m) navigation. Instrumentation is also available to determine temperatures to depths up to about 600 m below the seafloor during deep-sea drilling to establish the uniformity of heat flow to greater depths, and may be the only method for reliable measurements in shallow seafloor (o1 km) regions
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Figure 2 Diagrams of the three most commonly used marine heat flow probes. (Reproduced with permission from Louden KE and Wright JA (1989) Marine heat flow data: a new compilation of observations and brief review of its analysis. In: Wright JA and Londen KE (eds) Handbook of Seafloor Heat Flow, pp. 2–67, Boca Raton: CRC Press.)
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where ocean temperatures are more variable. Geothermal probes have also been developed for use with manned submersibles, allowing visual control of measurements in regions with large lateral gradients of heat flux or over specific geological features. A development is underway to install geothermal instrumentation on an autonomous underwater vehicle (ABE) for measurements in regions that are remote or difficult for normal ship operations. Environmental Corrections
Although most measurements on the deep seafloor with the usual instrumentation do not require corrections, they may be needed for some regions with unusual environmental parameters. As mentioned above, shallow water measurements may be subject to temporal bottom water temperature (BWT) variability, causing non-linear gradients in the seafloor. Using heat conduction theory, corrections may be applied if BWT variability is monitored for a sufficient period (months to years if possible) before the geothermal measurements. Conversely, nonlinear temperature–depth profiles can be inverted to yield BWT history assuming that the initial gradient was linear. This inverse procedure is non-unique, although closely spaced measured temperatures to a sufficient depth below the seafloor and well-determined thermal conductivity can reduce uncertainties. Continental margins and some deep-sea trenches are examples of regions where non-linear gradients have been measured, usually correlated with strong and variable deep currents that are probably focused by the seafloor topography. Topography causes lateral heat flow variability even with uniform and constant BWT, because the seafloor topography distorts isotherms that would otherwise be horizontal below a flat surface. Seafloor sedimentation reduces the heat flow measured because recently deposited sediments modify the seafloor boundary condition to which the equilibrium gradient must adjust. Corrections usually become significant when sedimentation rates exceed a few tens of meters per million years and/or the total sediment thickness exceeds 1 km. Vertical pore water flow in sediments may either enhance (for upward flow) or reduce (for downward flow) temperature gradients because the water advects some of the heat otherwise conducted upwards. The effects become significant for flow rates greater than a few centimeters per year, and nonlinear gradients may be expected if upward flow rates exceed about 10 cm year1. It is unusual for rates on normal deep seafloor to exceed the latter value because the fluid permeability of pelagic sediments is too low, but
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coarser and sometimes rapidly deposited sediments of continental margins may support relatively rapid flows.
Range of Measured Parameters As discussed later, the plate tectonic paradigm predicts that the highest thermal gradients and heat fluxes should be found in the youngest seafloor, and the lowest values in the oldest. Although this is generally true when measurements are averaged over regions of various seafloor ages, considerable variability occurs over the youngest seafloor as a result of vigorous hydrothermal circulation. The cooling of hot and permeable upper ocean crust supports seawater convection in the crust over lateral circulation scales of at least several to a few tens of kilometers; vertical scales probably do not exceed a few kilometers. The highest gradients and heat fluxes (up to 1001C m1 and 100 W m2, respectively) are measured in sediments near seafloor vents where fluids upwell, and the lowest (o0.0051C m1 and 0.005 W m2, respectively) where fluids downwell. Mean upper basement fluid velocities may be several meters per year, and much greater in conduits that support vigorous seafloor venting. The detailed pore water flow and permeability structure for any system of seafloor hydrothermal circulation have not yet been investigated in detail and are probably very complex. After the seafloor ages to 50–70 My, hydrothermal circulation appears to decrease to a level where modulation of the surface heat flux is small to negligible (Figure 3), probably dependent on the thickness and uniformity of sediment cover. For the oldest seafloor (100–180 My), heat flux is relatively uniform at about 50 mW m2710%. This value probably reflects the secular cooling of the upper oceanic mantle. Thermal conductivity of marine sediments generally varies less than a factor of two, from about 0.7 to o1.4 Wm1 K1. The lowest values are associated with red clays or siliceous oozes, which have up to 70% or more of water by weight and the highest with carbonate oozes or coarse sediments near continental margins with a high proportion of quartz. Since water has a low thermal conductivity (0.6 W m1 K1), sediments with a large percentage of water have low thermal conductivity. Since calcium carbonate and quartz have high thermal conductivity (about 3 W m1 K1 and 4 W m1 K1, respectively), sediments with a large percentage of either of these minerals have high thermal conductivity. Conductivity variations with depth may be caused
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climate (e.g., glacial) cycles. When thermal conductivity varies rapidly with depth the temperature–depth profile may not be linear. If thermal conductivity has been measured at closely spaced depth intervals, then the heat flow can be calculated using a method introduced by Bullard. This approach assumes the absence of significant heat sources or sinks and onedimensional, steady-state, conductive heat flow. Thus there is a linear relationship between temperature and the thermal resistance of the sediments (the sum from the surface to the depth of the temperature measurement of the inverse of thermal conductivity times the sediment thickness for that given thermal conductivity). The slope determined from a least-squares fit of this ‘Bullard plot’ gives the appropriate conductive heat flux.
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Figure 3 Data and models for ocean depth (A) and heat flow (B) as a function of age. (Reproduced with permission from Stein and Stein, 1992.) The data are averaged in 2 My bins, and one standard deviation about the mean value for each is shown by the envelope. Shown are the plate model of Parsons and Sclater (1977) (PSM), a cooling halfspace model with the same thermal parameters (HS), and the GDH1 plate model from Stein and Stein (1992). (C) Heat flow fraction (observed heat flow/GDH1 prediction) with age averaged in 2 My bins. The discrepancy for ages o50–70 My presumably indicates the fraction of the heat transported by hydrothermal flow. The fractions for ages o50 My (closed circles), which were not used in deriving GDH1, are fit by a least squares line. The sealing age, where the line reaches the fractional value of one, is 65710 My.
by: (1) turbidity flows that sort grain sizes into layers with different proportions of pore water; (2) ice rafting in higher latitudes that may intersperse higher conductivity sediments derived from the continents and lower conductivity marine pelagic sediments; and (3) variability in the proportion of calcium carbonate in sediments near the equator deposited over
Oceanic lithosphere forms at midocean ridges, where hot magma upwells, and then cools to form plates as the material moves away from the spreading center. As the plate cools, heat flow decreases and the seafloor deepens (Figure 3). However, only shallow (less than 1 km) measurements of lithospheric temperatures are possible. Hence, the two primary data sets used to constrain models for the variation in lithospheric temperature with age are seafloor depths and heat flow. The depth, corrected for sediment load, depends on the temperature integrated over the lithospheric thickness. The heat flow is proportional to the temperature gradient. Initially seafloor depths rapidly increase, with the average increase relative to the ridge crest proportional to the square root of the crustal age. However, for ages greater than about 50–70 My, the average increase in depth is slower and the curve is said to ‘flatten.’ Mean heat flow also decreases rapidly away from the ridge crest, with values approximately proportional to the inverse of the square-root of the age, but after about 50 My this curve also ‘flattens.’ Halfspace and Plate Models
Two different mathematical models are often used to describe the thermal evolution of the oceanic lithosphere, the halfspace (or boundary layer) and plate models. For the halfspace model, the predicted lithospheric thickness increases proportionally with the square root of age. Hence depth and heat flow vary as the square root of age and the reciprocal of the square root of age, respectively. However, the halfspace model cannot explain the ‘flattening’ of the curves. Alternatively, the plate model represents the
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lithosphere as a layer with a fixed constant temperature at its base. Initially, the cooling for the plate model is the same as the halfspace model, but for older ages the influence of the lower boundary results in slower cooling, approximately predicting the observed flattening of the depths and heat flow at older ages. The plate base is assumed to represent a depth at which additional heat is supplied from the mantle below to prevent the halfspace cooling at older ages. However, the model does not directly describe how this heat is added.
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understanding of the physics of processes affecting lithospheric thermal evolution improve, new and better reference models will be developed.
Application to Lithospheric Processes Using a reference model for the expected heat flow, regions can be examined to observe if the measured heat flow differs from that predicted and, if so, to study the causes of the discrepancy. Two primary discrepancies are assumed to reflect hydrothermal circulation and midplate swells.
Reference Models Hydrothermal Circulation
Heat flow measurements for crust of ages 0–65 My are generally lower than the predictions of all commonly used reference models (Figure 3). Because the heat flow measurements primarily reflect conductive heat flow, this difference has been attributed to hydrothermal water flow in the crust and sediments transporting some of the heat assumed in thermal models to be transferred by conduction. The missing heat transported by convection must appear somewhere, either as high conductive heat flow or as advective discharge to the sea. It is estimated that of the predicted global oceanic heat flux of 32 1012 W,
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Thermal models are solutions to the inverse problem of finding the temperature as a function of age that best fits the depth and heat flow. The data are used to estimate the primary model parameters (typically plate thickness, basal temperature, and thermal expansion coefficient), subject to other parameters generally specified a priori. The global average depth/heat flow variation with age, calculated for a specific model, data, and parameters is used as a reference model to represent ‘normal’ oceanic lithosphere. ‘Anomalies’, deviations from the reference model, are then investigated to see if they reflect a significant difference in thermal (or other) processes from the global average. It is important to realize that anomalies (defined as those observed minus the model-predicted values) for various reference models may differ significantly, and hence lead to different tectonic inferences. Until recently, a 125 km-thick plate model by Parsons and Sclater (denoted here PSM) was commonly used. Subsequently, as more data became available, it was noted that PSM systematically overpredicts depths and underpredicts heat flow for lithosphere older than 70–100 My, causing widespread ‘anomalies.’ A later joint inversion of the depth and heat flow data by Stein and Stein found that these ‘anomalies’ are reduced significantly by a plate model termed GDH1. GDH1 has a thinner lithosphere (95710 km), and a basal temperature of 145071001C, consistent with the PSM estimate (135072751C). GDH1 predicts heat flow in mW m2 as a function of age, t, in millions of years equal to 510 t1/2 for ages less than about 55 My, and 48 þ 96 exp( 0.0278 t) for older ages. Inversion of the same data, while prescribing a basal temperature of 13501C (a typically assumed temperature for upwelling magma at the ridge based on results from experimental petrology), also yields a thin (100 km) plate, with a somewhat higher value of lithospheric conductivity and hence makes very similar predictions. As the quality of the observed data and our
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Figure 4 Portion of the heat flow data from the FlankFlux survey on the Juan de Fuca Ridge sites where bare rock penetrates the sediment column have dramatically higher heat flow than the surrounding areas. Away from these outcrops, heat flow varies inversely with depth to basement rock. Expected heat flow for its age is shown with the dashed line. (Reproduced with permission from Davis EE, Chapman DS, Mottl MJ et al. (1992) FlankFlux: an experiment to study the nature of hydrothermal circulation in young oceanic crust. Canadian Journal of Earth Sciences 29: 925–952.)
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approximately one third occurs by hydrothermal flow. Characteristic heat flow patterns in young crust appear associated with water flow. Surveys at sites like the Galapagos Spreading Center show high heat flow associated with presumed upwelling zones at basement highs or fault scarps, and low heat flow associated with presumably down-flowing water at basement lows. Heat flow over the highs may be two to ten (or more) times greater than at nearby sites where water may be recharging the system (Figure 4). Closely spaced surveys in young crust show high scatter, presumably in part due to the local variations in sediment distribution, basement relief, and crustal permeability. The most spectacular evidence for hydrothermal circulation are vents near the ridge crest where hydrothermally altered water at temperatures up to 4001C exits the seafloor. Hot rock or magma at shallow depths provides a large heat source for the flow. The high sulfide content of the hot fluids supports unusual local biota from bacteria to tube worms. Vent areas show complex flow patterns with very high heat flow close to low heat flow. Although most vigorous vents are near spreading centers, on older crust isolated crystalline crust outcrops surrounded by a well-sedimented area can also vent measurable amounts of hydrothermal fluid (e.g., the FlankFlux survey area on the Juan de Fuca plate). The persistence of the heat flow discrepancy well away from ridges indicates that hydrothermal heat loss occurs in older crust. The ‘sealing age,’ defined as that beyond which measured heat flow approximately equals that predicted, is presumed to indicate the near cessation of heat transfer by hydrothermal circulation. The sealing age for the entire global data set is about 65710 My. Hence, although there are presumably local variations, water flow in older crust seems not, in general, to transport significant amounts of heat. However, some heat flow surveys (e.g., in the north-west Atlantic Ocean on 80 My crust and the Maderia Abyssal Plain on 90 My crust) indicate that hydrothermal circulation may continue, even if relatively little heat is lost by convection into the sea water. The reasons for the ‘sealing’ are not well understood. An earlier view was that 100– 200 m of sediment would be sufficiently impermeable to ‘seal’ off the crust from the sea, so that heat flow at these sites would yield a ‘reliable’ value, i.e., that predicted by conduction-only thermal models. However, many such sites have lower than expected values. A simple way to reconcile these observations with the present understanding of seafloor hydrology to assume that if water cannot flow vertically through thick sediments at a particular site, it may
flow laterally to a fault or basement outcrop, and then be manifested as either high conductive heat flow or hot water exiting to the sea, such as observed for the FlankFlux area. It has been suggested that the porosity and permeability of the crust decrease with increasing age, thus significantly reducing the water flow. However, recent studies of permeability and seismic velocity suggest that most of the rapid change occurs within the first 5–15 My and that older crust may still retain some relatively permeable pathways. Hydrothermal circulation has profound implications for the chemistry of the oceanic crust and sea water, because sea water reacts with the crust, giving rise to hydrothermal fluid of significantly different composition. The primary geochemical effects are thought to result from the high-temperature water flow observed at ridge axes. Circulation of sea water through hot rock removes magnesium and sulfate from sea water and enriches calcium, potassium, silica, iron, manganese, and other elements within the hydrothermal solution. Estimating the volume of water flowing through the crust depends on the heat capacity of the water, the heat lost by convection, and the assumed temperature of the water, the latter having the largest uncertainties. Although the heat flow anomaly is greatest in young lithosphere, only about 30% of the hydrothermal heat loss and 7% of the water flow occurs in crust younger than 1 My. This effect should be significant for ocean chemistry because a water volume equivalent to that of the total ocean is estimated to circulate through the oceanic crust about once every 0.5–5 My. Hot Spots
Oceanic midplate swells are identified by seafloor depths shallower than expected for their lithospheric age. Thus, models of the processes giving rise to these regions rely on assessments of how their heat flow and other properties differ from unperturbed lithosphere. The origin of these swells is generally thought to be related to upwelling mantle plumes (hot spots) that result in uplift and volcanism. Two basic types of models have been proposed for their origin. In one, the swell is a thermal effect due to the hot spot thinning and heating the lithosphere at depth. In the second, the uplift is primarily due to the dynamic effects of the upwelling plume, which may largely reflect thermal buoyancy forces within the upwelling mantle. The thermal models predict significantly larger heat flow anomalies than do the dynamic models. Detailed heat flow measurements have been made for Hawaii, Cape Verde, Reunion, Bermuda, and Crozet swells. Relative to the PSM reference curve, large heat flow anomalies are suggested.
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Figure 5 Heat flow data for transects along (lower left) and across (lower right) the Hawaiian Swell at the locations shown (upper left). The heat flow, though anomalously high with respect to the Parsons and Sclater (PSM) model, is at most slightly above that expected for GDH1. The predicted heat flow values are for 100 My (lower right) and 95–110 My (lower left). Figure modified from Von Herzen et al. (1982, 1989). (Reproduced with permission from Stein CA and Stein S (1993) Constraints of Pacific midplate swells from global depth-age and heat flow-age models. The Mesozoic Pacific, American Geophysical Union Monograph 77: 53–76.)
However, relative to both measured heat flow from off-swell lithosphere and the GDH1 reference curve, smaller heat flow anomalies (less than 5–8 mW m2) are deduced, supporting the idea of a primary dynamic mechanism for hotspots (Figure 5).
Application to Marine Margin Studies Subduction Zones
Many extensive marine heat flow surveys have been done near the Japanese subduction zones. On average, observed heat flow from the trench axis to the forearc area is lower than that characteristic of the crust seaward of the trench. However, heat flow is higher and more variable over the volcanic arc and back arc region compared to the area seaward of the trench. Recent surveys for subduction zones, including Barbados, Nankai, and Cascadia, show that heat flow is highly variable. Within accretionary prisms, high values are often associated with upward advection of pore fluids, typically found along faults and the bottom decollement. Although they have not been
investigated extensively, the active nonaccretionary forearcs seem to have the lowest mean heat fluxes (20– 30 mW m2). Relatively low heat flow values over the subducting lithosphere are the thermal consequence of the subduction of one plate beneath another. Overall, heat flow depends on the age, rate, geometry, and thermal structure of the subducting plate and the sediment thickness and deformation history of the region. Heat flow in marginal basins behind the volcanic arcs is often high. Many back arc basins with high heat flow appear to have formed by back-arc spreading processes similar to those at midocean ridges, within the last 50 My. Passive Continental Margins
Passive continental margins form after continental crust is rifted and seafloor spreading occurs. Because rifting heats the lithosphere, heat flow data can be used to constrain models of rifting. Subsequent to rifting, margins slowly subside as cooling occurs. Simple models for this process suggest that subsequent to rifting the additional heat will almost completely dissipate within 100 My. Although the
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amount of radioactive heating from continental crust may introduce some variability, older passive margins typically have heat flows similar to old oceanic crust, whereas young margins, such as in the Red Sea, have high heat flow.
Gas Hydrates
Gas hydrates, solid composites of biogenically derived gasses (mainly methane) combined with water ice, may be present in marine sediments, especially at continental margins, under the correct pressure and temperature conditions. Gas hydrates are detected by direct sampling, and inferred from seismic reflection data when a strong bottom-simulating reflector (BSR) is produced by the seismic velocity contrast between the gas hydrate and the sediment below. Given a depth (and thus pressure) of a bottom-simulating reflector, its temperature can be predicted from known relationships and a heat flow calculated. Alternatively, heat flow data can be used to estimate the bottom-simulating reflector temperature. The volume of hydrocarbons contained in marine gas hydrates is large. Changes in eustatic sea level (hence pressure) or bottom seawater temperatures could result in their release and thus increase greenhouse gasses and affect global climate.
Further Reading Humphris S, Mullineaux L, Zierenberg R, and Thomson R (eds.) (1995) Seafloor hydrothermal systems, physical, chemical, biological, and geological interactions, Geophysical Monographs, 91, 425--445. Hyndman RD, Langseth MG, and Von Herzen RP (1987) Deep Sea Drilling Project geothermal measurements: a review. Review of Geophysics 25: 1563--1582. Langseth MG, Jr and Von Herzen RP (1971) Heat Sow through the Soor of the world oceans. In: Maxwell AE (ed.) The Sea, vol. IV, part 1, pp. 299--352. New York: Wiley-Interscience. Lowell RP, Rona PA, and Von Herzen RP (1995) Seafloor hydrothermal systems. Journal of Geophysical Research 100: 327--352. Parsons B and Sclater JG (1977) An analysis of the variation of ocean floor bathymetry and heat flow with age. Journal of Geophysical Research 82: 803--827. Pollack HN, Hurter SJ, et al. (1993) Heat flow from the earth’s interior: analysis of the global data set. Review of Geophysics 31: 267--280. Stein CA and Stein S (1992) A model for the global variation in oceanic depth and heat flow with lithospheric age. Nature 359: 123--129. Von Herzen RP (1987) Measurement of oceanic heat flow. In: Sammis C and Henyey T (eds.) Methods of Experimental Physics – Geophysics, vol. 24, Part B, pp. 227--263. London: Academic Press. Wright JA and Louden KE (eds.) (1989) CRC Handbook of Seafloor Heat Flow. Boca Raton: CRC Press.
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES K. Lambeck, Australian National University, Canberra, ACT, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1157–1167, & 2001, Elsevier Ltd.
Introduction Geological, geomorphological, and instrumental records point to a complex and changing relation between land and sea surfaces. Elevated coral reefs or wave-cut rock platforms indicate that in some localities sea levels have been higher in the past, while observations elsewhere of submerged forests or flooded sites of human occupation attest to levels having been lower. Such observations are indicators of the relative change in the land and sea levels: raised shorelines are indicative of land having been uplifted or of the ocean volume having decreased, while submerged shorelines are a consequence of land subsidence or of an increase in ocean volume. A major scientific goal of sea-level studies is to separate out these two effects. A number of factors contribute to the instability of the land surfaces, including the tectonic upheavals of the crust emanating from the Earth’s interior and the planet’s inability to support large surface loads of ice or sediments without undergoing deformation. Factors contributing to the ocean volume changes include the removal or addition of water to the oceans as ice sheets wax and wane, as well as addition of water into the oceans from the Earth’s interior through volcanic activity. These various processes operate over a range of timescales and sea level fluctuations can be expected to fluctuate over geological time and are recorded as doing so. The study of such fluctuations is more than a scientific curiosity because its outcome impacts on a number of areas of research. Modern sea level change, for example, must be seen against this background of geologically–climatologically driven change before contributions arising from the actions of man can be securely evaluated. In geophysics, one outcome of the sea level analyses is an estimate of the viscosity of the Earth, a physical property that is essential in any quantification of internal convection and thermal processes. Glaciological modeling of the behavior of large ice sheets during the last cold
period is critically dependent on independent constraints on ice volume, and this can be extracted from the sea level information. Finally, as sea level rises and falls, so the shorelines advance and retreat. As major sea level changes have occurred during critical periods of human development, reconstructions of coastal regions are an important part in assessing the impact of changing sea levels on human movements and settlement.
Tectonics and Sea Level Change Major causes of land movements are the tectonic processes that have shaped the planet’s surface over geological time. Convection within the high-temperature viscous interior of the Earth results in stresses being generated in the upper, cold, and relatively rigid zone known as the lithosphere, a layer some 50–100 km thick that includes the crust. This convection drives plate tectonics — the movement of large parts of the lithosphere over the Earth’s surface — mountain building, volcanism, and earthquakes, all with concomitant vertical displacements of the crust and hence relative sea level changes. The geological record indicates that these processes have been occurring throughout much of the planet’s history. In the Andes, for example, Charles Darwin identified fossil seashells and petrified pine trees trapped in marine sediments at 4000 m elevation. In Papua New Guinea, 120 000-year-old coral reefs occur at elevations of up to 400 m above present sea level (Figure 1). One of the consequences of the global tectonic events is that the ocean basins are being continually reshaped as mid-ocean ridges form or as ocean floor collides with continents. The associated sea level changes are global but their timescale is long, of the order 107 to 108 years, and the rates are small, less than 0.01 mm per year. Figure 2 illustrates the global sea level curve inferred for the past 600 million years from sediment records on continental margins. The long-term trends of rising and falling sea levels on timescales of 50–100 million years are attributed to these major changes in the ocean basin configurations. Superimposed on this are smaller-amplitude and shorter-period oscillations that reflect more regional processes such as large-scale volcanism in an ocean environment or the collision of continents. More locally, land is pushed up or down episodically
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
and, in any discussion of global changes in sea level, information from such localities is best set aside in favor of observations from more stable environments.
Glacial Cycles and Sea Level Change
600 400 200 0 590
500
440 410
360 290 250 210 Time (million years BP)
140
65
35
5
Sea level above present level (m)
Pliocene_Pleistocene
Miocene
Paleocene & Eocene
Oligocene
Cretaceous
Permian
Carboniferous
Devonian
Silurian
Cambrian
Ordovician
in response to the deeper processes. The associated vertical crustal displacements are rapid, resulting in sea level rises or falls that may attain a few meters in amplitude, but which are followed by much longer periods of inactivity or even a relaxation of the original displacements. The raised reefs illustrated in Figure 1, for example, are the result of a large number of episodic uplift events each of typically a meter amplitude. Such displacements are mostly local phenomena, the Papua New Guinea example extending only for some 100–150 km of coastline. The episodic but local tectonic causes of the changing position between land and sea can usually be identified as such because of the associated seismic activity and other tell-tale geological signatures. The development of geophysical models to describe these local vertical movements is still in a state of infancy
Jurassic
Figure 1 Raised coral reefs from the Huon Peninsula, Papua New Guinea. In this section the highest reef indicated (point 1) is about 340 m above sea level and is dated at about 125 000 years old. Elsewhere this reef attains more than 400 m elevation. The top of the present sea cliffs (point 2) is about 7000 years old and lies at about 20 m above sea level. The intermediate reef tops formed at times when the rate of tectonic uplift was about equal to the rate of sea level rise, so that prolonged periods of reef growth were possible. Photograph by Y. Ota.
Triassic
More important for understanding sea level change on human timescales than the tectonic contributions – important in terms of rates of change and in terms of their globality – is the change in ocean volume driven by cyclic global changes in climate from glacial to interglacial conditions. In Quaternary time, about the last two million years, glacial and interglacial conditions have followed each other on timescales of the order of 104–105 years. During interglacials, climate conditions were similar to those of today and sea levels were within a few meters of their present day position. During the major glacials, such as 20 000 years ago, large ice sheets formed in the northern hemisphere and the Antarctic ice sheet expanded, taking enough water out of the oceans to lower sea levels by between 100 and 150 m. Figure 3 illustrates the changes in global sea level over the last 130 000 years, from the last interglacial, the last time that conditions were similar to those of today, through the Last Glacial Maximum and to the present. At the end of the last interglacial, at 120 000 years ago, climate began to fluctuate; increasingly colder conditions were reached, ice sheets over North America and Europe became more or less permanent features of the landscape, sea levels reached progressively lower values, and large parts of today’s coastal shelves were exposed. Soon after the culmination of maximum glaciation, the ice sheets again disappeared, within a period of about
_200
Figure 2 Global sea level variations through the last 600 million years estimated from seismic stratigraphic studies of sediments deposited at continental margins. Redrawn with permission from Hallam A (1984) Pre-Quaternary sea-level changes. Annual Review of Earth and Planetary Science 12: 205–243.
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
Relative sea level (m)
20 0 _20 _40 _60 _80 _100 _120
0
20
60 80 100 40 Time (thousand years BP)
120
Figure 3 Global sea level variations (relative to present) since the time of the last interglacial 120 000–130 000 years ago when climate and environmental conditions were last similar to those of the last few thousand years. Redrawn with permission from Chapell J et al. (1996) Reconciliation of Late Quaternary sea level changes derived from coral terraces at Huon Peninsula with deep sea oxygen isotope records. Earth and Planetary Science Letters 141b: 227–236.
10 000 years, and climate returned to interglacial conditions. The global changes illustrated in Figure 3 are only one part of the sea level signal because the actual ice– water mass exchange does not give rise to a spatially uniform response. Under a growing ice sheet, the Earth is stressed; the load stresses are transmitted through the lithosphere to the viscous underlying mantle, which begins to flow away from the stressed area and the crust subsides beneath the ice load. When the ice melts, the crust rebounds. Also, the meltwater added to the oceans loads the seafloor and the additional stresses are transmitted to the mantle, where they tend to dissipate by driving flow to unstressed regions below the continents. Hence the seafloor subsides while the interiors of the continents rise, causing a tilting of the continental margins. The combined adjustments to the changing ice and water loads are called the glacio- and hydro-isostatic effects and together they result in a complex pattern of spatial sea level change each time ice sheets wax and wane.
Observations of Sea Level Change Since the Last Glacial Maximum Evidence for the positions of past shorelines occurs in many forms. Submerged freshwater peats and tree stumps, tidal-dwelling mollusks, and archaeological sites would all point to a rise in relative sea level since the time of growth, deposition, or construction. Raised coral reefs, such as in Figure 1, whale bones cemented in beach deposits, wave-cut rock platforms
51
and notches, or peats formed from saline-loving plants would all be indicative of a falling sea level since the time of formation. To obtain useful sea level measurements from these data requires several steps: an understanding of the relationship between the feature’s elevation and mean sea level at the time of growth or deposition, a measurement of the height or depth with respect to present sea level, and a measurement of the age. All aspects of the observation present their own peculiar problems but over recent years a substantial body of observational evidence has been built up for sea level change since the last glaciation that ended at about 20 000 years ago. Some of this evidence is illustrated in Figure 4, which also indicates the very significant spatial variability that may occur even when, as is the case here, the evidence is from sites that are believed to be tectonically stable. In areas of former major glaciation, raised shorelines occur with ages that are progressively greater with increasing elevation and with the oldest shorelines corresponding to the time when the region first became ice-free. Two examples, from northern Sweden and Hudson Bay in Canada, respectively, are illustrated in Figure 4A. In both cases sea level has been falling from the time the area became ice-free and open to the sea. This occurred at about 9000 years ago in the former case and about 7000 years later in the second case. Sea level curves from localities just within the margins of the former ice sheet are illustrated in Figure 4B, from western Norway and Scotland, respectively. Here, immediately after the area became ice-free, the sea level fell but a time was reached when it again rose. A local maximum was reached at about 6000 years ago, after which the fall continued up until the present. Farther away from the former ice margins the observed sea level pattern changes dramatically, as is illustrated in Figure 4C. Here the sea level initially rose rapidly but the rate decreased for the last 7000 years up to the present. The two examples illustrated, from southern England and the Atlantic coast of the United States, are representative of tectonically stable localities that lie within a few thousand kilometers from the former centers of glaciation. Much farther away again from the former ice sheets the sea level signal undergoes a further small but significant change in that the present level was first reached at about 6000 years ago and then exceeded by a small amount before returning to its present value. The two examples illustrated in Figure 4D are from nearby localities in northern Australia. At both sites present sea level was reached about 6000 years ago after a prolonged period of rapid rise with resulting highstands that are small in amplitude but geographically variable.
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
Shoreline elevation (m)
250 140
Angermanland, northern Sweden
200
Ottawa Island, Hudson Bay, Canada
120 100
150 80 100
60 40
50 20 0 10
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50 Andoya, northern Norway
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_ 20
Dungeness, Kent,
8
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Cape Charles, Virginia, USA
0
_ 30 12
10
8
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Karumba, Gulf of Carpentaria, North Queensland
_ 10 2 _ 20
2
_ 30
0
1 −1 6 5 4 3 2 1 0 North Queensland, Australia
_ 40 _ 50 12
(D)
2
30
southern England (C)
3
Firth of Forth, Scotland
40
10
Shoreline elevation (m)
8
50
_ 10 20
(B)
0
10
8
6
1
0 0 6 5 4 2 4 Time (1000 radiocarbon years BP)
3
2
1
0
˚ ngermanland, Figure 4 Characteristic sea level curves observed in different localities (note the different scales used). (A) From A Gulf of Bothnia, Sweden, and from Hudson Bay, Canada. Both sites lie close to centers of former ice sheet, over northern Europe and North America respectively. (B) From the Atlantic coast of Norway and the west coast of Scotland. These sites are close to former icesheet margins at the time of the last glaciation. (C) From southern England and Virginia, USA. These sites lie outside the areas of former glaciation and at a distance from the margins where the crust is subsiding in response to the melting of the nearby icesheet. (D) Two sites in northern Australia, one for the Coral Sea coast of Northern Queensland and the second from the Gulf of Carpentaria some 200 km away but on the other side of the Cape York Peninsula. At both localities sea level rose until about 6000 years ago before peaking above present level.
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
The examples illustrated in Figure 4 indicate the rich spectrum of variation in sea level that occurred when the last large ice sheets melted. Earlier glacial cycles resulted in a similar spatial variability, but much of the record before the Last Glacial Maximum has been overwritten by the effects of the last deglaciation.
Glacio-Hydro-Isostatic Models If, during the decay of the ice sheets, the meltwater volume was distributed uniformly over the oceans, then the sea level change at time t would be Dze ðtÞ ¼ ðchange in ice volume=ocean surface areaÞ ri =rw
½1a
where ri and rw are the densities of ice and water, respectively. (See end of article – symbols used.) This term is the ice-volume equivalent sea level and it provides a measure of the change in ice volume through time. Because the ocean area changes with time a more precise definition is Dze ðtÞ ¼
ð dVi ri ri dt dt r0 A0 ðtÞ
½1b
t
where A0 is the area of the ocean surface, excluding areas covered by grounded ice. The sea level curve illustrated in Figure 3 is essentially this function. However, it represents only a zero-order approximation of the actual sea level change because of the changing gravitational field and deformation of the Earth. In the absence of winds or ocean currents, the ocean surface is of constant gravitational potential. A planet of a defined mass distribution has a family of such surfaces outside it, one of which – the geoid – corresponds to mean sea level. If the mass distribution on the surface (the ice and water) or in the interior (the load-forced mass redistribution in the mantle) changes, so will the gravity field, the geoid, and the sea level change. The ice sheet, for example, represents a large mass that exerts a gravitational pull on the ocean and, in the absence of other factors, sea level will rise in the vicinity of the ice sheet and fall farther away. At the same time, the Earth deforms under the changing load, with two consequences: the land surface with respect to which the level of the sea is measured is time-dependent, as is the gravitational attraction of the solid Earth and hence the geoid. The calculation of the change of sea level resulting from the growth or decay of ice sheets therefore
53
involves three steps: the calculation of the amount of water entering into the ocean and the distribution of this meltwater over the globe; the calculation of the deformation of the Earth’s surface; and the calculation of the change in the shape of the gravitational equipotential surfaces. In the absence of vertical tectonic motions of the Earth’s surface, the relative sea level change Dzðj; tÞ at a site j and time t can be written schematically as Dzðj; tÞ ¼ Dze ðtÞ þ DzI ðj; tÞ
½2
where DzI(j, t) is the combined perturbation from the uniform sea level rise term [1]. This is referred to as the isostatic contribution to relative sea level change. In a first approximation, the Earth’s response to a global force is that of an elastic outer spherical layer (the lithosphere) overlying a viscous or viscoelastic mantle that itself contains a fluid core. When subjected to an external force (e.g., gravity) or a surface load (e.g., an ice cap), the planet experiences an instantaneous elastic deformation followed by a timedependent or viscous response with a characteristic timescale(s) that is a function of the viscosity. Such behavior of the Earth is well documented by other geophysical observations: the gravitational attraction of the Sun and Moon raises tides in the solid Earth; ocean tides load the seafloor with a time-dependent water load to which the Earth’s surface responds by further deformation; atmospheric pressure fluctuations over the continents induce deformations in the solid Earth. The displacements, measured with precision scientific instruments, have both an elastic and a viscous component, with the latter becoming increasingly important as the duration of the load or force increases. These loads are much smaller than the ice and water loads associated with the major deglaciation, the half-daily ocean tide amplitudes being only 1% of the glacial sea level change, and they indicate that the Earth will respond to even small changes in the ice sheets and to small additions of meltwater into the oceans. The theory underpinning the formulation of planetary deformation by external forces or surface loads is well developed and has been tested against a range of different geophysical and geological observations. Essentially, the theory is one of formulating the response of the planet to a point load and then integrating this point-load solution over the load through time, calculating at each epoch the surface deformation and the shape of the equipotential surfaces, making sure that the meltwater is appropriately distributed into the oceans and that the total ice–water mass is preserved. Physical inputs into the
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
formulation are the time–space history of the ice sheets, a description of the ocean basins from which the water is extracted or into which it is added, and a description of the rheology, or response parameters, of the Earth. For the last requirement, the elastic properties of the Earth, as well as the density distribution with depth, have been determined from seismological and geodetic studies. Less well determined are the viscous properties of the mantle, and the usual procedure is to adopt a simple parametrization of the viscosity structure and to estimate the relevant parameters from analyses of the sea level change itself. The formulation is conveniently separated into two parts for schematic reasons: the glacio-isostatic effect representing the crustal and geoid displacements due to the ice load, and the hydro-isostatic effect due to the water load. Thus ½3 Dzðj; tÞ ¼ Dze ðtÞ þ Dzi ðj; tÞ þ Dzw ðj; tÞ where Dzi(j,t) is the glacio-isostatic and Dzw(j,t) the hydro-isostatic contribution to sea level change. (In reality the two are coupled and this is included in the formulation). If Dzi(j,t) is evaluated everywhere, the past water depths and land elevations H(j,t) measured with respect to coeval sea level are given by ½4 H ðj; tÞH0 ðjÞ Dzðj; tÞ where HðfÞ is the present water depth or land elevation at location j. A frequently encountered concept is eustatic sea level, which is the globally averaged sea level at any time t. Because of the deformation of the seafloor during and after the deglaciation, the isostatic term Dzðj; tÞ is not zero when averaged over the ocean at any time t, so that the eustatic sea level change is Dzeus ðtÞ ¼ Dze ðtÞ þ hDzI fj; tgi0 where the second term on the right-hand side denotes the spatially averaged isostatic term. Note that Dze ðtÞ relates directly to the ice volume, and not Dzeus ðtÞ.
The Anatomy of the Sea Level Function The relative importance of the two isostatic terms in eqn. [3] determines the spatial variability in the sea level signal. Consider an ice sheet of radius that is much larger than the lithospheric thickness: the limiting crustal deflection beneath the center of the load is Iri =Irm where I is the maximum ice thickness and Irm is the upper mantle density. This is the local isostatic approximation and it provides a reasonable approximation of the crustal deflection if the loading time is long compared with the relaxation time of the mantle. Thus for a 3 km thick ice sheet the maximum
deflection of the crust can reach 1 km, compared with a typical ice volume for a large ice sheet that raises sea level globally by 50–100 m. Near the centers of the formerly glaciated regions it is the crustal rebound that dominates and sea level falls with respect to the land. This is indeed observed, as illustrated in Figure 4A, and the sea level curve here consists of essentially the sum of two terms, the major glacio-isostatic term and the minor ice-volume equivalent sea level term Dze ðtÞ (Figure 5A). Of note is that the rebound continues long after all ice has disappeared, and this is evidence that the mantle response includes a memory of the earlier load. It is the decay time of this part of the curve that determines the mantle viscosity. As the ice margin is approached, the local ice thickness becomes less and the crustal rebound is reduced and at some stage is equal, but of opposite sign, to Dze(t). Hence sea level is constant for a period (Figure 5B) before rising again when the rebound becomes the minor term. After global melting has ceased, the dominant signal is the late stage of the crustal rebound and levels fall up to the present. Thus the oscillating sea level curves observed in areas such as Norway and Scotland (Figure 4B) are essentially the sum of two effects of similar magnitude but opposite sign. The early part of the observation contains information on earth rheology as well as on the local ice thickness and the globally averaged rate of addition of meltwater into the oceans. Furthermore, the secondary maximum is indicative of the timing of the end of global glaciation and the latter part of the record is indicative mainly of the mantle response. At the sites beyond the ice margins it is the meltwater term Dze(t) that is important, but it is not the sole factor. When the ice sheet builds up, mantle material flows away from the stressed mantle region and, because flow is confined, the crust around the periphery of the ice sheet is uplifted. When the ice sheet melts, subsidence of the crust occurs in a broad zone peripheral to the original ice sheet and at these locations the isostatic effect is one of an apparent subsidence of the crust. This is illustrated in Figure 5C. Thus, when the ocean volumes have stabilized, sea level continues to rise, further indicating that the planet responds viscously to the changing surface loads. The early part of the observational record (e.g., Figure 4C) is mostly indicative of the rate at which meltwater is added into the ocean, whereas the latter part is more indicative of mantle viscosity. In all of the examples considered so far it is the glacial-rebound that dominates the total isostatic adjustment, the hydro-isostatic term being present but comparatively small. Consider an addition of water that raises sea level by an amount D. The local
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
55
300
300
Relative sea level change (m)
Prediction
ice
200
100
100
0
0 ESL
_ 100 (A)
200
20
Angerman River, Gulf of Bothnia
15
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5
_ 100
0
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Relative sea level change (m)
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ice
100
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_ 50 ESL
_ 100 _ 150 20
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200
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Observation
Upper Firth of Forth, Scotland
15
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_ 100 _ 150
0
20
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20 water
0
0 Relative sea level change (m)
_ 20
ice
_ 20
_ 40
_ 40
_ 60
_ 60 _ 80
ESL
_ 80
_ 100
_ 100 _ 120
Dungeness, south of England
_ 140 (C)
15
20
Relative sea level change (m)
10
(D)
0 _ 10
10
5
20
10
15
5
0
0 _ 10
ice ESL
Karumba, Gulf of Carpentaria
_ 30 8
_ 140
10
water
_ 20 _ 40 10
0
_ 120
6
4
2
0
_ 20 _ 30 _ 40 10
8
6
4
2
0
Time (1000 years BP)
Figure 5 Schematic representation of the sea level curves in terms of the ice-volume equivalent function Dze ðtÞ (denoted by ESL) and the glacio- and hydro-isostatic contributions Dzi ðf; tÞ, Dzw ðf; tÞ (denoted by ice and water, respectively). The panels on the left indicate the predicted individual components and the panels on the right indicate the total predicted change compared with the observed values (data points). (A) For the sites at the ice center where Dzi ðf; tÞbDze ðtÞbDzw ðf; tÞ. (B) For sites near but within the ice margin where jDzi ðf; tÞjBjDze ðtÞjbjDzw ðf; tÞj, but the first two terms are of opposite sign. (C) For sites beyond the ice margin where jDze ðtÞj > jDzi ðf; tÞjbjDzw ðf; tÞj. (D) For sites at continental margins far from the former ice sheets where jDzw ðf; tÞj > jDzi ðtÞj. Adapted with permission from Lambeck K and Johnston P (1988). The viscosity of the mantle: evidence from analyses of glacial rebound phenomena. In: Jackson ISN (ed.) The Earth’s Mantle – Composition, Structure and Evolution, pp. 461–502. Cambridge: Cambridge University Press.
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
Sea Level Change and Shoreline Migration: Some Examples Figure 6 illustrates the ice-volume equivalent sea level function Dze ðtÞ since the time of the last maximum glaciation. This curve is based on sea level
indicators from a number of localities, all far from the former ice sheets, with corrections applied for the isostatic effects. The right-hand axis indicates the corresponding change in ice volume (from the relation [1b]). Much of this ice came from the ice sheets over North America and northern Europe but a not insubstantial part also originated from Antarctica. Much of the melting of these ice sheets occurred within 10 000 years, at times the rise in sea level exceeding 30 mm per year, and by 6000 years ago most of the deglaciation was completed. With this sea level function, individual ice sheet models, the formulation for the isostatic factors, and a knowledge of the topography and bathymetry of the world, it becomes possible to reconstruct the paleo shorelines using the relation [4]. Scandinavia is a well-studied area for glacial rebound and sea level change since the time the ice retreat began about 18 000 years ago. The observational evidence is quite plentiful and a good record of ice margin retreat exists. Figure 7 illustrates examples for two epochs. The first (Figure 7A), at 16 000 years ago, corresponds to a time after onset of deglaciation. A large ice sheet existed over
20
Time (radiocarbon, 1000 years BP) 15 10 5
0
0
0
_20
7.7
_40 _60
23.1
_80 _100
38.6
Ice volume (106 km3 )
isostatic response to this load is Drw =rm , where rw is the density of ocean water. This gives an upper limit to the amount of subsidence of the sea floor of about 30 m for a 100 m sea level rise. This is for the middle of large ocean basins and at the margins the response is about half as great. Thus the hydro-isostatic effect is significant and is the dominant perturbing term at margins far from the former ice sheets. This occurs at the Australian margin, for example, where the sea level signal is essentially determined by Dze ðtÞ and the water-load response (Figure 4D). Up to the end of melting, sea level is dominated by Dze ðtÞ but thereafter it is determined largely by the water-load term Dzw ðj; tÞ, such that small highstands develop at 6000 years. The amplitudes of these highstands turn out to be strongly dependent on the geometry of the waterload distribution around the site: for narrow gulfs, for example, their amplitude increases with distance from the coast and from the water load at rates that are particularly sensitive to the mantle viscosity. While the examples in Figure 5 explain the general characteristics of the global spatial variability of the sea level signal, they also indicate how observations of such variability are used to estimate the physical quantities that enter into the schematic model (3). Thus, observations near the center of the ice sheet partially constrain the mantle viscosity and central ice thickness. Observations from the ice sheet margin partially constrain both the viscosity and local ice thickness and establish the time of termination of global melting. Observations far from the ice margins determine the total volumes of ice that melted into the oceans as well as providing further constraints on the mantle response. By selecting data from different localities and time intervals, it is possible to estimate the various parameters that underpin the sea level eqn. [3] and to use these models to predict sea level and shoreline change for unobserved areas. Analyses of sea-level change from different regions of the world lead to estimates for the lithospheric thickness of between 60–100 km, average upper mantle viscosity (from the base of the lithosphere to a depth of 670 km) of about (1–5)1020 Pa s and an average lower mantle viscosity of about (1–5)1022 Pa s. Some evidence exists that these parameters vary spatially; lithosphere thickness and upper mantle viscosity being lower beneath oceans than beneath continents.
Ice-volume equivalent sea level (m)
56
_120 _140 25
54.0 5 20 15 10 Time (calibrated, 1000 years BP)
0
Figure 6 The ice-volume equivalent function Dze ðtÞ and ice volumes since the time of the last glacial maximum inferred from corals from Barbados, from sediment facies from north-western Australia, and from other sources for the last 7000 years. The actual sea level function lies at the upper limit defined by these observations (continuous line). The upper time scale corresponds to the radiocarbon timescale and the lower one is calibrated to calendar years. Adapted with permission from Fleming K et al. (1998) Refining the eustatic sea-level curve since the LGM using the far- and intermediate-field sites. Earth and Planetary Science Letters 163: 327–342; and Yokoyama Y et al. (2000) Timing of the Last Glacial Maximum from observed sea-level minimum. Nature 406: 713–716.
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
57
Figure 8 Shoreline reconstructions for South East Asia and northern Australia at the time of the last glacial maximum, (A) at about 18 000 years ago and (B) at 12 000 years ago. The water depth contours are the same as in Figure 7. The water depths in the inland lakes at 18 000 years ago are contoured at 25 m intervals relative to their maximum levels that can be attained without overflowing.
Figure 7 Shoreline reconstructions for Europe (A) at 16 000 and (B) at 10 500 years ago. The contours are at 400 m intervals for the ice thickness (white) and 25 m for the water depths less than 100 m. The orange, yellow, and red contours are the predicted lines of equal sea level change from the specified epoch to the present and indicate where shorelines of these epochs could be expected if conditions permitted their formation and preservation. The zero contour is in yellow; orange contours, at intervals of 100 m, are above present, and red contours, at 50 m intervals, are below present. At 10 500 years ago the Baltic is isolated from the Atlantic and its level lies about 25 m above that of the latter.
Scandinavia with a smaller one over the British Isles. Globally sea level was about 110 m lower than now and large parts of the present shallow seas were exposed, for example, the North Sea and the English Channel but also the coastal shelf farther south such as the northern Adriatic Sea. The red and orange contours indicate the sea level change between this period and the present. Beneath the ice these rebound contours are positive, indicating that if shorelines could form here they would be above sea level today. Immediately beyond the ice margin, a broad but shallow bulge develops in the topography, which will subside as the ice sheet retreats. At the second epoch selected (Figure 7B) 10 500 years ago, the ice has retreated and reached a temporary halt as the climate briefly returned to colder conditions; the Younger Dryas time of Europe. Much of the Baltic was then ice-free and a freshwater lake developed at some
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GLACIAL CRUSTAL REBOUND, SEA LEVELS, AND SHORELINES
25–30 m above coeval sea level. The flooding of the North Sea had begun in earnest. By 9000 years ago, most of the ice was gone and shorelines began to approach their present configuration. The sea level change around the Australian margin has also been examined in some detail and it has been possible to make detailed reconstructions of the shoreline evolution there. Figure 8 illustrates the reconstructions for northern Australia, the Indonesian islands, and the Malay–IndoChina peninsula. At the time of the Last Glacial Maximum much of the shallow shelves were exposed and deeper depressions within them, such as in the Gulf of Carpentaria or the Gulf of Thailand, would have been isolated from the open sea. Sediments in these depressions will sometimes retain signatures of these pre-marine conditions and such data provide important constraints on the models of sea level change. Part of the information illustrated in Figure 6, for example, comes from the shallow depression on the Northwest Shelf of Australia. By 12 000 years ago the sea has begun its encroachment of the shelves and the inland lakes were replaced by shallow seas.
Symbols used Ice-volume equivalent sea level. Uniform change in sea level produced by an ice volume that is distributed uniformly over the ocean surface. Perturbation in sea level due to glacioiDzI ðtÞ sostatic Dzi(t) and hydro-isostatic Dzw(t) effects. Dzeus ðtÞ Eustatic sea level change. The globally averaged sea level at time t. Dze ðtÞ
See also Beaches, Physical Processes Affecting. Coral Reefs. Fiordic Ecosystems. Geomorphology. Lagoons. Mangroves. Past Climate from Corals. Rocky Shores. Salt Marshes and Mud Flats. Salt Marsh Vegetation. Sandy Beaches, Biology of. Sea Level Change.
Further Reading Lambeck K (1988) Geophysical Geodesy: The Slow Deformations of the Earth. Oxford: Oxford University Press. Lambeck K and Johnston P (1999) The viscosity of the mantle: evidence from analysis of glacial-rebound phenomena. In: Jackson ISN (ed.) The Earth’s Mantle – Composition, Structure and Evolution, pp. 461--502. Cambridge: Cambridge University Press. Lambeck K, Smither C, and Johnston P (1998) Sea-level change, glacial rebound and mantle viscosity for northern Europe. Geophysical Journal International 134: 102--144. Peltier WR (1998) Postglacial variations in the level of the sea: implications for climate dynamics and solid-earth geophysics. Reviews in Geophysics 36: 603--689. Pirazzoli PA (1991) World Atlas of Holocene Sea-Level Changes. Amsterdam: Elsevier. van de Plassche O (ed.) (1986) Sea-Level Research: A Manual for the Collection and Evaluation of Data. Norwich: Geo Books. Sabadini R, Lambeck K, and Boschi E (1991) Glacial Isostasy, Sea Level and Mantle Rheology. Dordrecht: Kluwer.
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GLIDERS C. C. Eriksen, University of Washington, Seattle, WA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Underwater gliders are a recently developed class of autonomous underwater vehicle (AUV) driven by buoyancy changes to fly along saw-tooth trajectories through the ocean. Gliders quickly have become the AUVs with the highest endurance and longest range. They are able to sample the ocean interior at comparatively low cost because they can operate independently of ships, some for the better part of a year under global remote control. Sampling the ocean on space and timescales as fine as those dominating its variability and over long ranges and durations has been a challenge throughout the history of ocean science. Traditionally, most knowledge of the ocean interior has been collected by ships. Starting in the late twentieth century, moored observations have complemented those from ship-based surveys, and more recently satellite remote sensing has provided images of sea surface conditions globally on ever-finer space and timescales. Were these techniques inexpensive, there would be no particular need for autonomous data collection by moving platforms. Unfortunately, ships stay at sea only for a month or two and rarely are directed to sample the same region for multiple cruises. Moorings sampling the open ocean number in the hundreds globally, typically in the dozens in the deep sea. This means variability of the ocean interior is generally undersampled. Undersampling is arguably the principal impediment to description and understanding of the ocean. Gliders have been developed to address the shortcomings of conventional observing means so that such dominant sources of variability as mesoscale eddies, fronts, and boundary currents can be resolved simultaneously in space and time and over sufficiently long periods and wide domains to allow them to be understood. They are able to make progress in solving the sampling problem simply by the numbers their economy affords. The salient characteristic of gliders is that they cost roughly the equivalent of a few days of research vessel operation to build and can be operated for a few months for the cost of a day of ship time.
The tasks gliders do best are those for which ships are least suited: intensive, regular, and sustained observations of oceanic properties that are readily measured by electronic means. Observations that require large or power-hungry instruments, physical collection of samples, or specialized labor are unsuitable for gliders and continue the need for ship support. Gliders excel at measuring standard properties typically collected by a conductivity–temperature–depth (CTD) package. Among their many advantages is their ability to function well through the most severe seas the ocean has to offer, day and night, around the globe. They are a means to erasing the fair-weather bias of ship-based observations. One may sample the ocean at deliberately chosen locations and times with gliders, examine the data very shortly after it is collected, and alter the sampling plan as often as a glider reaches the sea surface and communicates. The density of observations can readily be scaled to the phenomena of interest by adjusting the number and distribution of platforms to address local, regional, or global variability on diurnal, fortnightly, or seasonal scales, for example. Of course, gliders are not without limitation: long range and high endurance are achieved at the cost of traveling slowly through the ocean. When they cease to communicate, they are lost.
A Short History John Swallow began the era of exploring the ocean interior autonomously in the 1950s with acoustically tracked floats ballasted to follow water motions at depth, famously deployed in the various North Atlantic locales. In following decades, acoustic floattracking technology extended to ocean basin scale. Russ Davis and Doug Webb turned to satellite tracking to follow the motion of Autonomous Lagrangian Circulation Explorer (ALACE) floats that periodically adjusted their buoyancy to reach the sea surface for communication, and then returned to a ‘parking’ depth to drift. ALACE floats were the precursor to Argo floats, platforms used in an ambitious international program to seed the ocean with 3000 profiling floats. While Argo floats continue to track horizontal water movements, the focus of the Argo program is largely on CTD profiles collected periodically to describe large-scale features of ocean circulation. Gliders are the descendants of profiling floats. They can be thought of as profiling floats with wings.
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Their essential difference is that their sampling location is controlled, subject to a limited ability to navigate through currents. Floats randomly sample where currents carry them. While gliders do not offer a Lagrangian view of the ocean, they are commanded to describe oceanic conditions at deliberate times and locations, adding to the Eulerian description of the ocean. Henry Stommel recognized the potential of underwater gliders (the idea was shared with him by his friend, neighbor, and colleague Doug Webb), and published a science fiction article describing them in 1989. Written as the retrospective of a scientist working at a glider network control station in 2021, the article has proved remarkably prescient in describing the gliders that were developed over the next decade and a half. Stommel envisioned a fleet of about 1000 gliders plying the oceans, continuously collecting profiles and communicating them ashore via satellite, while being controlled to variously occupy hydrographic sections or chase features within the ocean. The fundamental vision was of a distributed network of relatively inexpensive platforms. His optimism on the pace of development can understandably be blamed on his enthusiasm, for he predicted the maiden voyage at sea of a glider to have taken place in 1994 and last 198 days. In reality, the first ocean deployment longer than a week took place in 1999 and the longest mission as of this writing lasted 217 days, accomplished in 2005. The Autonomous Oceanographic Sampling Network program set up by the US Office of Naval Research (ONR) in 1995 provided 5 years of support that resulted in gliders becoming a reality. The three operational gliders, Slocum (Webb Research Corp.), Spray (Scripps Institution of Oceanography and Woods Hole Oceanographic Institution), and Seaglider (University of Washington), were developed under this program and have subsequently been supported both by ONR and the US National Science Foundation. Four key technical elements enabled the successful development of underwater gliders: small reliable buoyancy engines, low-power computer microprocessors, Global Positioning System (GPS) navigation, and low-power duplex satellite communication. Seaglider, Slocum, and Spray were designed to fulfill similar missions, hence understandably share many elements (Figure 1). All three are of similar size, roughly 50 kg in mass, in order that both manufacturing and operating costs would be relatively small. The vehicles can be launched and recovered from small vessels by two people without power-assisted equipment. They are designed to carry out missions from a few weeks (Slocum) to
several months (Seaglider and Spray) duration while traveling at B0.5 kt (B0.25 m s 1). All are batterypowered, navigate by dead-reckoning underwater between GPS navigational fixes obtained at the sea surface, and send data and receive commands via a constellation of low Earth-orbit satellites on longrange missions. Derivatives of these three gliders are currently under development, including a version powered in part by extracting thermal energy from ocean stratification, one that operates under ice, and one capable of making open ocean full-depth dives. In addition, parallel development of a vehicle considerably larger, faster, and capable of flying along a shallower glider path is underway.
Design Considerations The three operational gliders were designed to address the sampling deficiencies of infrequent shipbased surveys and widely separated moorings in describing oceanic internal structure. Mesoscale eddies with characteristic space and timescales of O(100 km) and O(1 month) provide the principal noise to ocean general circulation and climate variability. Fronts, eddies, internal waves, diurnal cycles, and tidal fluctuations contribute to variability on smaller scales in problems of regional focus. Measurement platforms capable of resolving these features require sampling on scales O(10 km) and O(10 h) while the need to observe many realizations of random processes calls for sampling to be carried out for several weeks or months over tens or hundreds of kilometers. In order to observe at much lower cost than with ships or moorings, the three operational gliders were designed with endurance and range capability to meet these demands. Gliding is the conversion of a buoyancy force into a body’s forward as well as vertical motion. While aeronautical gliders are heavier than the air around them, hence always fall through it, underwater gliders are designed to alternately be heavier and lighter than the surrounding ocean so that they may glide both when diving and climbing. On average, underwater gliders displace an amount of water equal to their mass. They accomplish gliding by adjusting their volume displacement downward to dive and upward to climb. Gliders pitch themselves downward and are supported by wings to descend along a slanting path (conversely to ascend). Wings are glider propellers. They convert the vertical force of buoyancy into forward motion. The desire to obtain vertical as well as horizontal structure of the ocean is conveniently and efficiently
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Figure 1 The operational gliders (clockwise from left): Seaglider pitched down at the sea surface with its trailing antenna raised (photo by Keith Van Thiel, University of Washington), top view of Spray at the sea surface (photo by Robert Todd, Scripps Institution of Oceanography), and rear view of Slocum climbing (photo by David Fratantoni, Woods Hole Oceanographic Institution).
GLIDERS
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provided by gliding, provided glide slopes are steep compared to slopes of oceanic property surfaces. The operational gliders variously operate along glide slopes in the range 0.25 to unity (e.g., B14–451), similar to the glide slope of a NASA space shuttle. By contrast, high-performance sail planes attain slopes as gentle as 0.02, not far from the higher slopes of property surfaces within the ocean. Sinking through the ocean is gratis, while climbing through it requires energy input. The force of ambient pressure can be used to partially collapse vehicle volume, raise its mass density over that of seawater surrounding it, and make it sink. Diving requires only the energy to open a valve to allow a swim bladder to drain into a reservoir within the hull. To stop diving and begin ascent, displacement volume must be increased, requiring work against pressure to expand a swim bladder, for example. Work is required not only to provide sufficient potential energy to bring a glider to the sea surface, but also to provide kinetic energy to fly through the ocean. Once made buoyant, gliding eventually brings a vehicle to the sea surface where it can raise an antenna to navigate and communicate. Kinetic energy provided through buoyancy forcing is ultimately dissipated by hydrodynamic drag in gliding. Since drag is generally quadratic with speed, halving speed quadruples endurance and doubles range. Underwater gliders achieve endurance and range 2 orders of magnitude higher than conventional propeller-driven AUVs by virtue of employing speeds an order of magnitude smaller. While a slow propeller-driven AUV might attain considerable endurance and range traveling horizontally, it too must work against the ocean’s potential energy gradient (gravity) to reach the sea surface to navigate or communicate electromagnetically. Slocum, Spray, and Seaglider were each designed to operate at speeds barely faster than typical ocean currents. Since currents are typically surface-intensified due to baroclinicity, deeper dives tend to reduce average current encountered by a glider, hence enhance the ability to navigate effectively. Gliders can navigate more effectively in regions where currents reverse with depth, despite being swift, than in regions of strong barotropic flow. For example, glider navigation in a swift equatorial current system or within a strait with strongly baroclinic flow is expected to be more effective than in weakly stratified coastal jets over the continental shelf. Gliders have operated across the Kuroshio and Gulf Stream by gliding with an upstream component to ferry across them, much as a canoeist paddles both upstream and across stream in order to cross a swift river. The finite spatial extent of ocean currents can be exploited to
make up for downstream drift within them by gliding upstream in weak or opposing flows on their flanks. Navigating the swirling currents of a strong eddy is also possible if glider position within an eddy is recognized. The operational gliders all carry a fixed energy supply in the form of primary batteries. Packs consisting of lithium cells provide the highest energy density, but are classified as hazardous material. Gliders using these are subject to various shipping restrictions, particularly as air freight. Nevertheless, lithium battery packs enable glider missions of half a year or more. Up to about one quarter of the total mass of these gliders can be devoted to carrying batteries, so packs of the equivalent of B100 lithium D-cells can be used, providing roughly 10 MJ of energy, enough for B1000 dive cycles using B10 kJ per cycle. The total average power consumption may be as little as 0.5 W, resulting in mission endurance of several months. Recent improvements suggest that missions longer than a year are feasible for some battery-powered gliders. A buoyancy engine powered by ocean thermal stratification, as envisioned by Stommel and Webb, is under development for Slocum, but will be restricted to regions where sufficient ocean temperature stratification is found. In principle, longer endurance and range can be attained by larger vehicles, since drag is determined by area while energy storage scales with volume. Instrument payload also scales with volume, and also expense. Practically, AUVs larger than B60 kg require power assistance for handling at sea. Both the significant economy of manual handling by a pair of operators and the smaller unit manufacture and service costs associated with small platforms are consistent with the vision expressed by Stommel’s 1989 article and also that guiding the Argo project: that of a network of a large number of relatively small, inexpensive platforms. By keeping gliders small, networks of them are tolerant to occasional vehicle loss.
Glider Dive Cycles Gliders are fully autonomous from the time they leave the sea surface until they return. At the surface, they use a GPS fix to compute a desired heading to a target (or use a preset or communicated heading), then reduce volume displacement to start diving. Retracting a piston accomplishes buoyancy reduction in the 200-m-depth rated Slocum, while in Spray and Seaglider an external bladder is deflated to reduce buoyancy by opening a valve, since subbarometric pressure is maintained inside the pressure
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GLIDERS
hull. Glider pitch attitude is adjusted by redistributing mass within the vehicles, both by changing buoyancy and moving fixed masses such as battery packs. Gliders turn to attain the desired heading through the use of an active rudder or by redistributing mass to roll the wings. Typical turn radii of the operational gliders are tens of meters. Under water, navigation is by dead reckoning: simply following the chosen compass course. While course through the water varies due to environmental effects such as water-column shear of horizontal current as well as to vehicle control, the overall horizontal displacement due to gliding at various headings and speeds can be measured or inferred. The difference between displacement between surface GPS fixes and gliding displacement gives an estimate of depth-averaged current. The depthaveraged current for each dive cycle can be used to guide the choice of vehicle heading to approach a target through a prediction scheme, either on board or communicated from afar.
Hydrodynamics The existing operational underwater gliders attain effectively steady flight in some tens of seconds due to relatively slow acceleration relative to hydrodynamic drag. The balance of buoyancy by lift and drag forces determines the glide slope. Glider design requires sufficiently small drag that the ratio of drag to lift, equal in steady flight to glide slope, can provide the desired performance. The wings of Seaglider, Slocum, and Spray are somewhat small relative to body vehicle size in comparison to those of typical aeronautical sail planes. Both considerations of how steeply to dive through the ocean’s structure and ease of handling on launch and recovery guided wing sizing. Vehicle drag depends on both shape and surface texture, form, and skin drag, respectively. A significant component of drag is that induced by lift, typically parametrized as proportional to the square of attack angle. Considerations of drag led Spray to approximate a slender ellipsoidal form with a long cylinder capped by ellipsoidal ends and Seaglider to use a long smooth forebody (to maintain laminar boundary layer flow) and short after-section (where turbulence is tolerated). Tow tank, wind tunnel, and open-ocean experience have demonstrated that small appendages can account for disproportionately large contributions to total vehicle drag. A glider with a given set of lift, drag, and induced drag coefficients can be flown over a range of speed and glide slope, the latter limited by stall at the
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maximum lift to drag ratio. As an example, buoyancy, power, and attack angle dependence on vertical and horizontal speed are shown in Figure 2 for a typical Seaglider. Contours of power (solid curves) reflect the nearly quadratic nature of the drag law for this vehicle: operating at speeds of B0.2 m s 1 costs B0.1 W, while operating at B0.4 m s 1 costs B0.4 W (of delivered power – the buoyancy engines are at most about 50% efficient). The maximum horizontal speed component attainable for a given power use is at a glide slope only slightly steeper than the stall glide slope, as can be seen from the departure of the power curve shape from near-circular about the origin to steeply curved to form a cone of exclusion at the stall slope. The extreme marks the most efficient glide slope for range, B0.28 for speeds of B0.25 m s 1 to B0.20 for speeds of B1 m s 1 for the example given in Figure 2. A given buoyancy force leads to a wide range of speed, depending on glide slope, as illustrated by the dashed curves in Figure 2. For example, a buoyancy of 150 g can provide a horizontal speed of 0.22 m s 1 near the glide slope of maximum efficiency, or 0.33 m s 1 at a glide slope of 0.8 (using about 4 times as much power). The maximum total speed is highest for a given buoyancy if a glider were to be able to rise perfectly vertically (requiring no wing lift), but the maximum horizontal speed is attained at slopes of 0.78–0.84 (glide angles of 30–401 above horizontal) for the particular combination of lift, drag, and induced drag coefficients in the Seaglider example. The buoyancy range plotted in Figure 2 exceeds what the Seaglider buoyancy engine is capable of providing in both positive and negative buoyancy (the end-to-end range is B850 g), but diving through a light near surface layer may produce instantaneously buoyancy as high as is plotted in the figure. Attack angles for gliders (the difference between pitch and glide angle, where the latter is always steeper) range from less than 11 for steep dive slopes to several degrees for glide slopes near stall. For an underwater flying wing, for example, glider pitch can be in the opposite sense to glide angle to produce large lift. For the Seaglider example, typical attack angles are a few degrees for the most efficient glide slope. Attack angle affects flow around the hull, amplifying it as a function of position.
Instrumentation Size, power, hydrodynamic smoothness, and mass invariance are the principal constraints on instrumentation carried on gliders. For the operational
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gliders, sensors and their packaging are practically limited to perhaps B500-cc volume and power consumption a fraction of that used for glider propulsion. On long missions, as much as 85% of glider battery energy might be devoted to running the buoyancy engine, limiting the average power consumption available to instrumentation and operating the onboard microprocessor to less than B0.1 W. Average power of this level is achieved by lowering the duty cycle of typical instruments considerably; that is, most of the time sensors are turned off and only briefly are energized for sampling. The low-power sleep current of the microprocessor is a significant constraint on mission endurance. Oceanic fields that are readily measured electronically are well suited to gliders. Temperature, salinity, and pressure are readily measured electronically at little energy cost. In addition to scientific
interest, together with measurements of vehicle pitch, volume, and heading, these can be used to estimate glider speed through the water. Other instrumentation routinely installed on gliders include dissolved oxygen sensors, bio-optical sensors for chlorophyll fluorescence, optical backscatter, and transmissivity, and acoustic sensors for both active and passive measurements. Specially adapted acoustic Doppler current profilers have also been carried on gliders. All of these instruments involve tradeoffs between capability, sampling rate, and mission endurance.
Missions The track of a Seaglider in the Labrador Sea (Figure 3) illustrates the ability of a glider to make transects in regions where depth-averaged flow is
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Figure 3 Track of a Seaglider in the Labrador Sea from 24 Sep. 2004 to 29 Apr. 2005, the longest AUV mission to date. This vehicle was launched from R/V Knorr in Davis Strait and recovered more than 7 months later offshore Nuuk, Greenland. Pink symbols mark the location of GPS fixes and Iridium calls between dive cycles. Blue arrows indicate depth-averaged currents (to the shallower of bottom depth and 1 km), where the scale length is typical glider speed through the water. Depth contour intervals are 100–1000 m, 500 m thereafter with heavy contours every 1 km.
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sufficiently weak, but be hampered in its ability to navigate within strong currents. Launched from a research vessel offshore west Greenland, this vehicle made two transects across Davis Strait before being sent south to Hamilton Bank offshore southeast Labrador. It was swept across part of the continental shelf region before escaping the Labrador Current to continue northeastward on its return to Greenland near Kap Desolation. On its way north toward recovery, it was caught in an anticyclonic eddy that drifted offshore. This glider was directed mainly northward as it cycled around the eddy core four times before it finally escaped by being guided radially outward across the eddy. It continued northeastward to Fyllas Bank, offshore Nuuk, Greenland, where it was recovered from a fishing vessel, having completed the longest AUV mission to date as of this writing, lasting 217 days, traveling 3750 km through the water, and making 663 dives, the majority to 1000-m depth. Several different types of glider missions have been undertaken, of which the example in Figure 3 is just one: a solitary glider survey. In other applications, gliders have been used to make repeat surveys along the same track, controlled to maintain position (a ‘virtual mooring’), used in numbers to survey a region intensively, used to interact with one another, to communicate with moored sensors, act as a courier for data recorded by these devices, and to control them.
See also Drifters and Floats. Underwater Vehicles.
Platforms:
Autonomous
Further Reading Davis RE, Eriksen CC, and Jones CP (2003) Autonomous buoyancy-driven underwater gliders. In: Griffiths G (ed.) Technology and Applications of Autonomous Underwater Vehicles, pp. 37--58. New York: Taylor and Francis. Rudnick DL, Davis RE, Eriksen CC, Fratantoni DM, and Perry MJ (2004) Underwater gliders for ocean research. Journal of the Marine Technology Society 38: 73--84. Stommel H (1989) The Slocum mission. Oceanography 2: 22--25.
Relevant Websites http://iop.apl.washington.edu – Operations Summary: Custom View, Seaglider Status, Integrative Observational Platforms, Ocean Physics Department, Applied Physics Laboratory, University of Washington. http://spray.ucsd.edu – SIO IDG Spray Home. http://www.webbresearch.com – Slocum Glider, Webb Research Corporation. http://seaglider.washington.edu – The Seaglider Fabrication Center of the University of Washington.
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GLOBAL MARINE POLLUTION A. D. Mclntyre, University of Aberdeen, Aberdeen, UK & 2009 Elsevier Ltd. All rights reserved.
This aritcle is designed to provide an entry to the historical and global context of issues related to pollution of the oceans. Until recently, the size and mobility of the oceans encouraged the view that they could not be significantly affected by human activities. Freshwater lakes and rivers had been degraded for centuries by effluents, particularly sewage, but, although from the 1920s coastal oil pollution from shipping discharges was widespread, it was felt that in general the open sea could safely dilute and disperse anything added to it. Erosion of this view began in the 1950s, when fallout from the testing of nuclear weapons in the atmosphere resulted in enhanced levels of artificial radionuclides throughout the world’s oceans. At about the same time, the effluent from a factory at Minamata in Japan caused illness and deaths in the human population from consumption of mercurycontaminated fish, focusing global attention on the potential dangers of toxic metals. In the early 1960s, buildup in the marine environment of residues from synthetic organic pesticides poisoned top predators such as fish-eating birds, and in 1967 the first wreck of a supertanker, the Torrey Canyon, highlighted the threat of oil from shipping accidents, as distinct from operational discharges. It might therefore be said that the decades of the 1950s and 1960s saw the beginnings of marine pollution as a serious concern, and one that demanded widespread control. It attracted the efforts of national and international agencies, not least those of the United Nations. The fear of effects of radioactivity focused early attention, and initiated the establishment of the International Commission on Radiological Protection (ICRP), which produced a set of radiation protection standards, applicable not just to fallout from weapons testing but also to the increasingly more relevant issues of operational discharges from nuclear reactors and reprocessing plants, from disposal of low-level radioactive material from a variety of sources including research and medicine, and from accidents in industrial installations and nuclear-powered ships. Following Minamata, other metals, in particular cadmium and lead, joined mercury on the list of concerns. Since this was seen, like radioactive wastes, as a public health problem, immediate action was taken. Metals
in seafoods were monitored and import regulations were put in place. As a result, metal toxicity in seafoods is no longer a major issue, and since most marine organisms are resilient to metals, this form of pollution affects ecosystems only when metals are in very high concentrations, such as where mine tailings reach the sea. Synthetic organics, either as pesticides and antifoulants (notably tributyl tin (TBT)) or as industrial chemicals, are present in seawater, biota, and sediments, and affect the whole spectrum of marine life, from primary producers to mammals and birds. The more persistent and toxic compounds are now banned or restricted, but since many are resistant to degradation and tend to attach to particles, the seabed sediment acts as a sink, from which they may be recirculated into the water column. Other synthetic compounds include plastics, and the increasing use of these has brought new problems to wildlife and amenities. Oil contaminates the marine environment mainly from shipping and offshore oil production activities. Major incidents can release large quantities of oil over short periods, causing immense local damage; but in the longer term, more oil reaches the sea via operational discharges from ships. These and other threats to the ocean are controlled by the International Convention for the Prevention of Pollution from Ships, administered by the International Maritime Organisation of the UN. Pollution is generated by human activities, and the most ubiquitous item is sewage, which is derived from a variety of sources: as a direct discharge; as a component of urban wastewater; or as sludge to be disposed of after treatment. Sewage in coastal waters is primarily a public health problem, exposing recreational users to pathogens from the local population. The dangers are widely recognized, and many countries have introduced protective legislation. At the global level, the London Dumping Convention controls, among other things, the disposal of sewage sludge. As well as introducing pathogens, sewage also contributes carbon and nutrients to the sea, adding to the substantial quantities of these substances reaching the marine environment from agricultural runoff and industrial effluents. The resulting eutrophication is causing major ecosystem impacts around the world, resulting in excessive, and sometimes harmful, algal blooms. The need for a global approach to ocean affairs was formally brought to the attention of the United
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Nations in 1967, and over the next 15 years, while sectoral treaties and agreements were being introduced, negotiations for a comprehensive regime led in 1982 to the adoption of the UN Convention on the Law of the Sea (UNCLOS). This provided a framework for the protection and management of marine resources. In the ensuing years, concepts such as sustainable development and the precautionary approach came to the fore and were endorsed in the Declaration of the Rio Summit, while proposals for Integrated Coastal Zone Management (ICZM) are being widely explored nationally and advanced through the Intergovernmental Oceanographic Commission. Early ideas of pollution were focused on chemical inputs, but following the Group of Experts on the Scientific Aspects of Marine Pollution (GESAMP) definition pollution is now seen in a much wider context, encompassing any human activity that damages habitats and amenities and interferes with legitimate users of the sea. Thus, manipulation of terrestrial hydrological cycles, and other hinterland activities including alterations in agriculture, or afforestation, can profoundly influence estuarine regimes, and are seen in the context of pollution. In particular, it is now recognized that excessive fishing can do more widespread damage to marine ecosystems than most chemical pollution, and the need for ecosystem-based fisheries management is widely recognized. Over the years, assessments of pollution effects have altered the priority of concerns, which today are very different from those of the 1950s. Thanks to the rigorous control of radioactivity and metals, these are not now major worries. Also, decades of experience with oil spills have shown that, after the initial damage, oil degrades and the resilience of natural communities leads to their recovery. Today, while the effects of sewage, eutrophication, and harmful algal blooms top the list of pollution concerns, along with the physical destruction of habitats by coastal construction, another item has been added: aquatic invasive species. The particular focus is on the transfer of harmful organisms in ships’ ballast water and sediments, which can cause disruption of fisheries, fouling of coastal industry, and reduction of human amenity. In conclusion, most of the impacts referred to above are on the shallow waters and the shelf, associated with continental inputs and activities. The open ocean, although subject to contamination from the atmosphere and from vessels in shipping lanes, is relatively less polluted United Nations Environment Programme (UNEP).
See also Anthropogenic Trace Elements in the Ocean. Antifouling Materials. Atmospheric Input of Pollutants. Chlorinated Hydrocarbons. Eutrophication. Exotic Species, Introduction of. International Organizations. Marine Policy Overview. Metal Pollution. Oil Pollution. Phytoplankton Blooms. Pollution: Approaches to Pollution Control. Pollution: Effects on Marine Communities. Pollution, Solids. Radioactive Wastes. Thermal Discharges and Pollution. Viral and Bacterial Contamination of Beaches.
Further Reading Brackley P (ed.) (1990) World Guide to Environmental Issues and Organisations. Harlow: Longman. Brune D, Chapman DV, and Gwynne DW (1997) Eutrophication. In: The Global Environment, ch. 30. Weinheim: VCH. Coe JM and Rogers DB (1997) Marine Debris. Berlin: Springer. de Mora SJ (ed.) (1996) Tributyltin Case Study of an Environmental Contaminant. Cambridge, UK: Cambridge University Press. GESAMP (1982) Reports and Studies No. 15: The Review of the Health of the Oceans. Paris: UNESCO. Grubb M, Koch M, and Thomson K (1993) The ‘Earth Summit’ Agreements: A Guide and Assessment. London: Earth Scan Publications. HMSO (1981) Eighth Report of the Royal Commission on Environmental Pollution, Oil Pollution of the Sea, Cmnd 8358. London: HMSO. Hollingworth C (2000) Ecosystem effects of fishing. ICES Journal of Marine Science 57(3): 465--465(1). International Maritime Organization (1991) IMO MARPOL 73/78 Consolidated Edition MARPOL 73/78. London: International Maritime Organization. International Oceanographic Commission (1998) Annual Report of the International Oceanographic Commission. Paris: UNESCO. Kutsuna M (ed.) (1986) Minamata Disease. Kunamoto: Kunammoyo University. Matheickal J and Raaymakers S (eds.) (2004) Second International Ballast Water Treatment R & D Symposium, 21–23 July 2003: Proceedings. GloBallast Monograph Series No. 15. London: International Maritime Organization. National Research Council (1985) Oil in the Sea. Washington, DC: National Academies Press. Park PK, Kester DR, and Duedall IW (eds.) (1983) Radioactive Wastes and the Ocean. New York: Wiley. Pravdic V (1981) GESAMP the First Dozen Years. Nairobi: UNEP. Pritchard SZ (1987) Oil Pollution Control. Wolfeboro, NH: Croom Helm. Sinclair M and Valdimarsson G (eds.) (2003) Responsible Fisheries in the Marine Ecosystem. Oxford, UK: CABI.
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GLOBAL MARINE POLLUTION
Tolba MK, El-Kholy OA, and El-Hinnawi E (1992) The World Environment 1972–1992: Two Decades of Challenge. London: Chapman and Hall. UNEP (1990) The State of the Marine Environment. UNEP Regional Seas Reports and Studies No. 115. Nairobi: UNEP.
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UNO (1983) The Law of the Sea: Official Text of UNCLOS. New York: United Nations. World Commission on Environment and Development (1987) Our Common Future. Oxford, UK: Oxford University Press.
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GRABS FOR SHELF BENTHIC SAMPLING P. F. Kingston, Heriot-Watt University, Edinburgh, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction The sedimentary environment is theoretically one of the easiest to sample quantitatively and one of the most convenient ways to secure such samples is by means of grabs. Grab samplers are used for both faunal samples, when the grab contents are retained in their entirety and then sieved to remove the biota from the sediment, and for chemical/physical samples when a subsample is usually taken from the surface of the sediment obtained. In both cases, the sampling program is reliant on the grab sampler taking consistent and relatively undisturbed sediment samples.
Conventional Grab Samplers The forerunner of the grab samplers used today is the Petersen grab, designed by C.G.J. Petersen to conduct benthic faunal investigations in Danish fiords in the early part of the twentieth century. It consisted of two quadrant buckets that were held in an open
Figure 1 Petersen grab.
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position and lowered to the seabed (Figure 1). On the bottom, the relaxing of the tension on the lowering warp released the buckets and subsequent hauling caused them to close before they left the bottom. The instrument is still used today but is seriously limited in its range of usefulness, working efficiently only in very soft mud. Petersen’s grab formed the basis for the design of many that came after. One enduring example is the van Veen grab, a sampler that is in common use today (Figure 2). The main improvement over Petersen’s design is the provision of long arms attached to the buckets to provide additional leverage to the closing action. The arms also provided a means by which the complex closing mechanism of the Petersen grab could be simplified with the hauling warp being attached to chains on the ends of the arms. The mechanical advantage of the long arms can be improved further by using an endless warp rig; this has the added advantage of helping to prevent the grab being jerked off the bottom if the ship rolls as the grab is closing. The van Veen grab was designed in 1933 and is still widely used in benthic infaunal studies owing to its simple design, robustness, and digging efficiency. The van Veen grab
Figure 2 Van Veen grab.
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Figure 3 Diagram of a Hunter grab. Reproduced from Hunter B and Simpson AE (1976) A benthic grab designed for easy operation and durability. Journal of the Marine Biological Association 56: 951–957.
typically covers a surface area of 0.1 m2, although instruments of twice this size are sometimes used. A more recent design of frameless grab is the Hunter grab (Figure 3). This is of a more compact design than the van Veen. The jaws are closed by levers attached to the buckets in a parallelogram configuration, giving the mechanism a good overall mechanical advantage. The closing action requires no chains or pulleys and the instrument can be operated by one person. Its disadvantage is that the bucket design does not encourage good initial penetration of the sediment, which is important in hardpacked sediments. A disadvantage of the grab samplers discussed so far is that there is little latitude for horizontal movement of the ship while the sample is being secured: the smallest amount of drift and the sampler is likely to be pulled over. The Smith–McIntyre grab was designed to reduce this problem by mounting the grab buckets in a stabilizing frame (Figure 4). Initial penetration of the leading edge of the buckets is assisted by the use of powerful springs and the buckets are closed by cables pulling on attached short arms in a similar way to that on the van Veen grab. The driving springs are released by two trigger plates, one on either side of the supporting frame to ensure that the sampler is resting flat on the seabed before the sample is taken. In firm sand the Smith–McIntyre grab penetrates to about the same depth of sediment as the van Veen grab. Its main disadvantage is the need to cock the spring mechanism on deck before deploying the sampler, a process that can be quite hazardous in rough weather. The Day grab is a simplified form of the Smith– McIntyre instrument in which the trigger and closing mechanism remains the same, but without spring assistance for initial penetration of the buckets
Figure 4 Smith–McIntyre grab.
Figure 5 Day grab.
(Figure 5). The Day grab is widely used, particularly for monitoring work, despite its poor performance in hard-packed sandy sediments. Most of the grabs thus far discussed have been designed to take samples with a surface area of 0.1 or 0.2 m2. The Baird grab, however, takes samples of 0.5 m2 by means of two inclined digging plates that are pulled together by tension on the warp (Figure 6). The grab is useful where a relatively large surface area needs to be covered, but has the disadvantage of taking a shallow bite and having the surface of the sample exposed while it is being hauled in.
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Warp Activation All the grabs described above use the warp acting against the weight of the sampler to close the jaws. However, direct contact with the vessel on the surface during the closure of the grab mechanism poses several problems.
Warp Heave
As tension is taken up by the warp to close the jaws, there is a tendency for the grab to be pulled up off the bottom, resulting in a shallower bite than might be expected from the geometry of the sampler. This tendency is related to the total weight of the sampler and the speed of hauling and is exacerbated by firm sediments. For example, the theoretical maximum depth of bite of a 120-kg Day grab is 13 cm (based
on direct measurements of the sampler); however, in medium sand, the digging performance is reduced to a maximum depth of only 8 cm (Figure 7(e)). The influence of warp action on the digging efficiency of a grab sampler can also depend on the way in which the sampler is rigged. This is particularly true of the van Veen grab. Figure 7(b) shows the bite profile of the chain-rigged sampler in which the end of each arm is directly connected to the warp by a chain. The vertical sides of the profile represent the initial penetration of the grab while the central rise shows the upward movement of the grab as the jaws close. Figure 7(c) shows the bite profile of a van Veen of similar size and weight (30 kg) rigged with an endless warp in which the arms are closed by a loop of wire passing through a block at the end of each arm (as in Figure 2). The vertical profile of the initial penetration is again apparent; however, in this case, the overall depth of the sampler in the sediment is maintained as the jaws close. The endless warp rig increases the mechanical advantage of the pull of the warp while decreasing the speed at which the jaws are closed. The result is that the sampler is ‘insulated’ from surface conditions to a greater extent than when chain-rigged, giving a better digging efficiency. Grab ‘Bounce’
In calm sea conditions it is relatively easy to control the rate of warp heave and obtain at least some consistency in the volume of sediment secured. However, such conditions are seldom experienced in the open sea where it is more usual to encounter wave action. Few ships used in offshore benthic
Figure 6 Baird grab.
(a)
(b)
(c)
2 4 6 8 10
2 4 6 8 10
2 4 6 8 10
Petersen grab (30 kg)
Chain-rigged van Veen grab (30 kg)
Endless-warp-rigged van Veen grab (short-armed, 30 kg)
(d)
(e)
(f)
2 4 6 8 10 12 14
2 4 6 8 10
2 4 6 8 10
Endless-warp-rigged van Veen grab (long-armed, 70 kg)
Day grab (120 kg)
Smith−mcintyre grab (120 kg)
Figure 7 Digging profiles of a range of commonly used benthic grab samplers obtained in a test tank using a fine sand substratum. (a) Peterson grab (30 kg); (b) chain-rigged van Veen grab (30 kg); (c) endless-warp-rigged van Veen grab (short-armed, 30 kg); (d) endless-warp-rigged van Veen grab (long-armed, 70 kg); (e) Day grab (120 kg); (f) Smith–McIntyre grab (120 kg).
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studies are fitted with winches with heave compensators so that the effect of ship’s roll is to introduce an erratic motion to the warp. This may result in the grab ‘bouncing’ off the bottom where the ship rises just as bottom contact is made, or in the grab being snatched off the bottom where the ship rises just as hauling commences. In the former instance, it is unlikely that any sediment is secured; in the latter, the amount of material and its integrity as a sample will vary considerably, depending on the exact circumstances of its retrieval. The intensity of this effect will depend on the severity of the weather conditions. Figure 8 shows the relationship between wind speed and grab failure rate, which is over 60% of hauls at wind force 8. What is of more concern to the scientist attempting to obtain quantitative samples is the dramatic increase in variability with increase in wind speed with a coefficient of variation between 20 and 30 at force 7. The high cost of ship-time places considerable pressure on operators to work in as severe weather conditions as possible and it is not unusual for sampling to continue in wind force 7 conditions with all its disadvantages. Drift
For a warp-activated grab sampler to operate efficiently it should be hauled with the warp positioned vertically above. Where there is a strong wind or current, these conditions may be difficult to achieve. The result is that the grab samplers are pulled on to their sides. This is a particular problem with samplers,
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such as the van Veen grab, that do not have stabilizing frames. Diver observations have shown, however, that at least in shallow water, where the drift effect is at its greatest on the bottom, even the framed heavily weighted Day and Smith–McIntyre grabs can be toppled. Initial Penetration
It is clear that the weight of the sampler is an important element in determining the volume of the sample secured. Much of the improved digging efficiency of the van Veen grab shown in Figure 7(d) can be attributed to the addition of an extra 40 kg of weight which increased the initial penetration of the sampler on contact with the sediment surface. Initial penetration is one of the most important factors in the sequence of events in grab operation, determining the final volume of sediment secured. Figure 9 shows the relationship between initial penetration and final sample volume obtained for a van Veen grab. Over 70% of the final volume is determined by the initial penetration. Subsequent digging of the sampler is hampered, as already shown, by the pull of the warp. For most benthic faunal studies it is important for the sampler to penetrate at least 5 cm into the sediment (for a 0.1 m2 surface area sample this gives 5 l of sample). In terms of number of species and individuals, over 90% of benthic macrofauna are found in the top 4–5 cm of sediment. Figure 10 shows how the number of individuals relates to average sample 20 18
70
16 Initial penetration depth (cm)
Number of failures due to motion as the % of number of attempts
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60 50 40 30 20
14 12 10 8 6 4
10
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0
0 0
1
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3
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0
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Sea state (Beaufort scale) Figure 8 Relationship between wind speed and grab failure rate.
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8
10 12 14 16 18 20 22 24
Sample volume (l) Figure 9 Relationship between initial penetration of a van Veen grab sampler and volume of sediment secured.
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1500
Number of individuals per 0.5 m2
1400 1300 1200 1100 1000 900 800 700 600
Figure 11 Shipek grab.
500 4.5
5
5.5
6
6.5
7
7.5
Average volume of grab sample (l) Figure 10 Relationship between number of benthic fauna individuals captured and sample volume for a boreal offshore sand substratum.
volume for 18 stations in the southern North Sea (each liter recorded represents 1 cm of penetration). Although there is a considerable variation in the numbers of individuals between stations, there is no significant trend linking increased abundance with increased sample volume penetration. No sample volumes of less than 4.5 l were taken, indicating that at that level of penetration most of the fauna were being captured. Samplers in which the jaws are held rigidly in a frame have no initial penetration if the edges of the jaw buckets, when held in the open position, are on a level with the base of the frame. The lack of any initial penetration in such instruments has the added disadvantage in benthic fauna work of undersampling at the edges of the bite profile (see Figures 7(e) and 7(f)), although the addition of weight will usually increase the sample volume obtained. Pressure Wave Effect
The descent of the grab necessarily creates a bow wave. Under field conditions, it is usually impracticable to lower the grab at a rate that will eliminate a preceding bow wave, even if the sea were flat calm. There have been several investigations of the effects of ‘downwash’ both theoretical, using artificially placed surface objects, and in situ. The effects of downwash can be reduced by replacing the upper surface of the buckets with an open mesh. Although there is still a considerable effect on the surface flock layer (rendering the samples of dubious
value for chemical contamination studies), the effect on the numbers of benthic fauna is generally very small. Self-Activated Bottom Samplers
There can be little doubt that one of the most important factors responsible for sampler failure or sample variability in heavy seas is the reliance of most presently used instruments on warp-activated closure. The most immediate and obvious answer to this problem is to make the closing action independent of the warp by incorporating a selfpowering mechanism. Spring-Powered Samplers
One solution to the problem is to use a spring to actuate the sampler buckets. Such instruments are in existence, possibly the most widely used being the Shipek grab, a small sampler (0.04 m2) consisting of a spring-loaded scoop (Figure 11). This instrument is widely used where small superficial sediment samples are required for physical or chemical analysis. The use of a pretensioned spring unfortunately sets practical limits on the size of the sampler, since to cock a spring in order to operate a sampler capable of taking a 0.1 m2 sample would require a force that would be impracticable to apply routinely on deck. In addition, in rough weather conditions, a loaded sampler of this size would be very hazardous to deploy. Compressed-Air-Powered Samplers
Another approach has been to use compressed air power. In the 1960s, Flury fitted a compressed air ram to a modified Petersen grab with success. However, the restricted depth range of the instrument and the inconvenience of having to recharge the
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GRABS FOR SHELF BENTHIC SAMPLING
air reservoir for each haul limited its potential for routine offshore work. Piston
Hydraulically Powered Samplers
Hydraulically powered grabs are commonly used for large-scale sediment shifting operations such as seabed dredging. The Bedford Institute of Oceanography, Nova Scotia, successfully scaled down this technology to that of a practical benthic sampler. Their instrument is relatively large, standing 2.5 m high and weighing some 1136 kg. It covers a surface area of 0.5 m2 and samples to a maximum sediment depth of 25 cm. At full penetration, the sediment volume taken is about 100 l. The buckets are driven closed by hydraulic rams powered from the surface. The grab is also fitted with an underwater television camera which allows the operator to visually select the precise sampling area on the seabed, close and open the bucket remotely, and verify that the bucket closed properly prior to recovery. The top of the buckets remain open during descent to minimize the effect of downwash and close on retrieval to reduce washout of the sample on ascent. The current operating depth of the instrument is 500 m. The instrument has been successfully used on several major offshore studies, but does require the use of a substantial vessel for its deployment. Hydrostatically Powered Samplers
Hydrostatically powered samplers use the potential energy of the difference in hydrostatic pressure at the sea surface and the seabed. The idea of using this power source is not new. In the early part of the twentieth century, a ‘hydraulic engine’ was in use by marine geologists that harnessed hydrostatic pressure to drive a rock drill. Hydrostatic power has also been used to drive corers largely for geological studies. However, these instruments were principally concerned with deep sediment corers and were not designed to collect macrofauna or material at the sediment–water interface. A more recent development has been that of a grab built by Heriot-Watt University, Edinburgh. The sampler uses water pressure difference to operate a hydraulic ram that is activated when the grab reaches the seabed. Figure 12 shows the general layout of the instrument. Water enters the upper chamber of the cylinder when the sampler is on the seabed, forcing down a piston that is connected to a system of levers that close the jaws. The actuating valve is held shut by the weight of the sampler and there is a delay mechanism to prevent premature closure of the jaws resulting from ‘bounce’. Back on the ship, the jaws are held shut by an overcenter locking mechanism
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Water inlet valve with release delay mechanism held shut by weight or sampler
Ram operated by hydrostatic pressure Connecting rod
Figure 12 Diagram of a grab sampler using hydrostatic pressure to close the jaws.
and, on release, are drawn open by reversal of the piston motion from air pressure built up on the underside of the piston during its initial power stroke. Since the powering of the grab jaws is independent of the warp, the sampler may be used successfully in a much wider range of surface weather conditions than conventional grabs.
Alternatives to Grab Samplers Ideally a benthic sediment sample for faunal studies should be straight-sided to the maximum depth of its excavation and should retain the original stratification of the sediment. Grab samplers by the very nature of their action will never achieve this end. Suction Samplers
One answer to this problem is to employ some sort of corer designed to take samples of sufficient surface area to satisfy the present approaches to benthic studies. The Knudsen sampler is such a device and is theoretically capable of taking the perfect benthic sample. It uses a suction technique to drive a core tube of 0.1 m2 cross-sectional area 30 cm into the sediment. Water is pumped out of the core tube on the seabed by a pump that is powered by unwinding a cable from a drum. The sample is retrieved by pulling the core out sideways using a wishbone arrangement and returning it to the surface bottomside up (Figures 13 and 14). Under ideal conditions, the device will take a straight-sided sample to a depth
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Figure 13 Knudsen sampler in descent position.
of 30 cm. However, conditions have to be flat calm in order to allow time for the pump to operate on the seabed and evacuate the water from the core. This limits the use of the Knudsen sampler and it is generally not suitable for sampling in unsheltered conditions offshore. Mounting the sampler in a stabilizing frame can improve its success rate and it is used regularly for inshore monitoring work where it is necessary to capture deep burrowing species. Spade Box Samplers
Another approach to the problem is to drive an open-ended box into the sediment, using the weight of the sampler, and arrange for a shutter to close off the bottom end. The most widespread design of such an instrument is that of the spade box sampler, first described by Reineck in the 1950s and later subjected to various modifications. The sampler consists of a removable steel box open at both ends and driven into the sediment by its own weight. The lower end of the box is closed by a shutter supported on an arm pivoted in such a way as to cause it to slide through the sediment and across the mouth of the box (Figure 15). As with the grab samplers previously described, the shutter is driven by the act of hauling on the warp with all the attendant disadvantages. Nevertheless, box corers are very successful and are used widely for obtaining relatively undisturbed samples of up to 0.25 m2 surface area (Figure 16). One big advantage of the box sampler is that the box can usually be removed with the sample and its overlying water left intact, allowing detailed studies of the sediment surface. Furthermore, it is possible to subsample using small-diameter corers for studies of chemical and physical characteristics. Despite their potential of securing the ‘ideal’ sediment sample, box corers are rarely used for routine benthic monitoring work. This is largely because of their size (a box corer capable of taking a 0.1 m2
Figure 14 Knudsen sampler in ascent position.
Lead weights
Box closure (‘spade’) Box holder Box
Figure 15 Diagram of a Reineck spade box sampler.
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Cable drum
Scoops Figure 17 Diagram of a Holme scoop. Reproduced from Holme NA (1953) The biomass of the bottom fauna in the English Channel off Plymouth. Journal of the Marine Biological Association 32: 1–49.
macrobenthos sampling. At present, there is no instrument that fulfils the requirements for a quick turnaround precision multiple corer for offshore sampling.
Sampling Difficult Sediments
Figure 16 A 0.25 m2 spade box sampler.
sample weighs over 750 kg and stands 2 m high) and the difficulty in deployment and recovery in heavy seas. Precision Corers
For chemical monitoring, it is important that the sediment–water interface is maintained intact, for it is the surface flock layer that will contain the most recently deposited material. Unfortunately, such undisturbed samples are rarely obtained using grab samplers or box corers. Precision corers are capable of securing undisturbed surface sediment cores; however, they are unsuitable for routine offshore work because of the time taken to secure a sample on the seabed and dependence on warp-activated closure. Additionally, the cross-sectional area of the core (0.002–0.004 m2) would necessitate the taking of large numbers of replicate samples in order to capture sufficient numbers of benthic macrofauna to be useful. This would be impracticable given the time taken to take a single sample. Large multiple precision corers have been constructed; these are usually too large and difficult to deploy for routine
Most of the samplers so far discussed operate reasonably well in mud or sand substrata. Few operate satisfactorily in gravel or stony mixed ground either because the bottom is too hard for the sampler to penetrate the substratum or because of the increased likelihood of a stone holding the jaws open when they are drawn together. To get around this problem various types of scoops have been devised. The Holme grab has a double scoop action with two buckets rotating in opposite directions to minimize any lateral movement during digging. The scoops are closed by means of a cable and pulley arrangement (Figure 17) and simultaneously take two samples of 0.05 m2 surface area. The Hamon grab, which has proved to be very effective in coarse, loose sediments, takes a single rectangular scoop of the substratum covering a surface area of about 0.29 m2. The scoop is forced into the sediment by a long lever driven by pulleys that are powered by the pull of the warp (Figure 18). Although the samples may not always be as consistent as those from a more conventional grab sampler, the Hamon grab has found widespread use where regular sampling on rough ground is impossible by any other means.
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A generic disadvantage of most samplers presently in use is that they rely on slackening of the warp to trigger the action and the heave of the warp to drive the closing mechanism. In calm conditions, this presents no great problem, but with increasing sea state the vertical movement of the warp decreases reliability dramatically until the variability and finally failure rate of hauls make further sampling effort fruitless.
Specific Problems, Requirements, and Future Developments Scoop
Release Stop plate hook Figure 18 Diagram of a Hamon scoop.
Present State of Technology It is perhaps surprising that given the high state of technology of survey vessels, position fixing, and analytical equipment, the most commonly used samplers are relatively primitive (being designed some 40 or more years ago). Yet the quality of the sample is of fundamental importance to any research or monitoring work. Currently the most popular instruments are grab samplers, probably because of their wide operational weather window and apparent reliability. Samplers such as the Day, van Veen, and Smith–McIntyre grab samplers are still routinely used for sampling sediment for chemical and biological analysis despite their well-documented shortcomings. As discussed earlier, the most important of these are the substantial downwash that precedes the sampler as it descends and the disturbance of the trapped sediment layers by the closing action of the jaws. Both chemical and meiofaunal studies are particularly vulnerable to these. Although box corers go some way to reducing disturbance of the sediment strata, the all-important surface flocculent layer is invariably washed away. A big disadvantage of the box corer for routine offshore work is that it is sensitive to weather conditions; in addition, the larger instruments do not perform well on sand substrata.
Chemical Studies
Studies involving sediment chemistry require precision sampling if undisturbed samples at the sediment/water interface are to be obtained. This can be critical, particularly when the results of recent sedimentation are of interest. The impracticality of using existing hydraulically damped corers such as the Craib corer for offshore work has led to the widespread use of less ‘weather sensitive’ devices such as spade box corers and grabs for routine monitoring purposes. However, studies carried out have shown that these samplers produce a considerable ‘downwash effect’, blowing the surface flocculent layer away before the sample is secured. This can have serious consequences if any meaningful estimation of the surface chemistry of the sediment is desired. Repetitive and accurate sampling is also a prerequisite for determining spatial and temporal change in sediment chemistry.
Meiofauna Studies Meiofauna has increasingly been shown to have potential as an important tool in benthic monitoring. One of the major factors limiting its wider adoption is the lack of a suitable sampler. Although instruments such as the Craib corer and its multicorer derivatives are capable of sampling the critically important superficial sediment layer, these designs provide a poor level of success on harder sediments and in anything but near-perfect weather conditions. They also have a slow turnaround time. Box corers are widely used as an alternative; however, they are known to be unreliable in their sampling of meiofauna. As with the macrobenthos, meiobenthic patchiness results in low levels of precision of abundance estimates unless large numbers of samples are taken.
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GRABS FOR SHELF BENTHIC SAMPLING
Macrobenthic Studies
The measurement and prediction of spatial and temporal variation in natural populations are of great importance to population biologists, both for fundamental research into population dynamics and productivity and in the characterization of benthic communities for determining change induced by environmental impact. Though always an important consideration, cost-effectiveness of sampling and sample processing is not so crucial in fundamental research, since time and funding may be tailored to fit objectives. This is rarely the case in routine monitoring work where often resolution and timescales have to fit the resources available. Benthic fauna are contagiously distributed and to sample such communities with a precision that will enable distinction between temporal variation and incipient change resulting from pollution effects, it is generally accepted that five replicate 0.1 m2 hauls from each station are necessary (giving a precision at which the standard error is no more than 20% of the mean). This frequency of sampling requires approximately 10–15 man-days of sediment faunal analysis per sample station. While this may be acceptable in community structure studies in which time and manpower (and thus cost) are not a primary consideration, this high cost of analyzing samples is of importance in routine monitoring studies and has led monitoring agencies and offshore operators to reduce sampling frequency on cost grounds to as few as two replicates per station. This reduces the precision with which faunal abundance can be estimated to a level at which only gross change can be demonstrated. However, sampling to an acceptable precision may be achieved from an area equivalent to 0.1 m2 if a smaller sampling unit is used. For example, 50 5-cm core samples (with a similar total surface area) have been shown to give a similar precision to that of 5 to 12 0.1-m2 grab samples. Thus a similar degree of precision may be obtained for around one-fifth to one-twelfth the analytical costs using a conventional approach. The problem is to be able to secure the 50 core samples per site that would be needed for the macrofaunal monitoring in a timescale that would be realistic offshore. At present, there is no instrument capable of supporting such a sampling demand and operating in the range of sediment types that wide-
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scale monitoring studies demand. Clearly a single core sampler would be impracticable, and one must look to the future development of a multiple corer that is capable of flexibility in its operation, which allows a quick on-deck turnaround between hauls.
See also Benthic Boundary Organisms Overview.
Layer
Effects.
Benthic
Further Reading Ankar S (1977) Digging profile and penetration of the van Veen grab in different sediment types. Contributions from the Asko¨ Laboratory, University of Stockholm, Sweden 16: 1--12. Beukema JJ (1974) The efficiency of the van Veen grab compared with the Reineck box sampler. Journal du Conseil Permanent International pour l’Exploration de la Mer 35: 319--327. Eleftheriou A and McIntyre AD (eds.) (2005) Methods for the Study of the Marine Benthos. Oxford, UK: Blackwell. Flury JA (1967) Modified Petersen grab. Journal of the Fisheries Research Board of Canada 20: 1549--1550. Holme NA (1953) The biomass of the bottom fauna in the English Channel off Plymouth. Journal of the Marine Biological Association 32: 1--49. Hunter B and Simpson AE (1976) A benthic grab designed for easy operation and durability. Journal of the Marine Biological Association 56: 951--957. Riddle MJ (1988) Bite profiles of some benthic grab samplers. Estuarine, Coastal and Shelf Science 29(3): 285--292. Thorsen G (1957) Sampling the benthos. In: Hedgepeth JW (ed.) Treatise on Marine Ecology and Paleoecology, Vol. 1: Ecology. Washington, DC: The Geological Society of America.
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GRAVITY M. McNutt, MBARI, Moss Landing, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1177–1185, & 2001, Elsevier Ltd.
Introduction The gravity field varies over the oceans on account of lateral variations in density beneath the ocean surface. The most prominent anomalies arise from undulations on density interfaces, such as occur at the water–rock interface at the seafloor or at the crust– mantle interface, also known as the Moho discontinuity. Because marine gravity is relatively easy to measure, it serves as a remote sensing tool for exploring the earth beneath the oceans. The interpretation of marine gravity anomalies in terms of the Earth’s structure is highly nonunique, however, and thus requires simultaneous consideration of other geophysically observed quantities. The most useful auxiliary measurements include depth of the ocean from echo sounders, the shape of buried reflectors from marine seismic reflection data, and/or the density of ocean rocks as determined from dredge samples or inferred from seismic velocities. Depending on the spatial wavelength of the observed variation in the gravity field, marine gravity observations are applied to the solution of a number of important problems in earth structure and dynamics. At the very longest wavelengths of 1000 to 10 000 km, the marine gravity field is usually combined with anomalies over land to infer the dynamics of the entire planet. At medium wavelengths of several tens to hundreds of kilometers, the gravity field contains important information on the thermal and mechanical properties of the lithospheric plates and on the thickness of their sedimentary cover. At even shorter wavelengths, the field reflects local irregularities in density, such as produced by seafloor bathymetric features, magma chambers, and buried ore bodies. On account of the large number of potential contributions to the marine gravity field, modern methods of analysis include spectral filtering to remove signals outside of the waveband of interest and interpretation within the context of models that obey the laws of phyics.
Units Gravity is an acceleration. The acceleration of gravity on the earth’s surface is about 9.81 m s2.
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Gravity anomalies (observed gravitational acceleration minus an expected value) are typically much smaller, about 0.5% of the total field. The SI-compatible unit for gravity anomaly is the gravity unit (gu): 1 gu ¼ 106 m s2. However, the older c.g.s unit for gravity anomaly, the milligal (mGal), is still in very wide use: 1 mGal ¼ 10 gu. Typical small-scale variations in gravity over the ocean range from a few tens to a few hundreds of gravity units. Lateral variations in gravitational acceleration (gravity gradients) are measured in Eo¨tvo¨s units (E): 1 E ¼ 109 s2. Another quantity useful in gravity interpretation is the density of earth materials, measured in kg m3. In the marine realm, relevant densities range from about 1000 kg m3 for water to more than 3300 kg m3 for mantle rocks. A close relative of the marine gravity field is the marine geoid. The geoid, measured in units of height, is the elevation of the sea level equipotential surface. Geoid anomalies are measured in meters and are the departure of the true equipotential surface from that predicted for an idealized spheroidal Earth whose density structure varies only with radius. Geoid anomalies range from 0 to more than 7100 m. The direction of the force of gravity is everywhere perpendicular to the geoid surface, and the magnitude of the gravitational attraction is the vertical derivative of the geopotential U (eqn [1]). g¼
@U @z
½1
Geoid height N is related to the same equipotential U via Brun’s formula eqn ([2]). N¼
U g0
½2
in which g0 is the acceleration of gravity on the spheroid. For a ship sailing on the sea surface (the equipotential), it is easier to measure gravity. From a satellite in free-fall orbit high over the Earth’s surface, radar altimeters can measure with centimeter precision variations in the elevation of sea level, an excellent approximation to the true geoid that would follow the surface of a motionless ocean. Regardless of whether geoid or gravity is the quantity measured directly, simple formulas in the wavenumber domain allow gravity to be computed from geoid and vice versa. Given the same equipotential, the gravity representation emphasizes the power in
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GRAVITY
the high-frequency (short-wavelength) part of the spectrum, whereas the geoid representation emphasizes the longer wavelengths. Therefore, for investigations of high-frequency phenomena, gravity is generally the quantity interpreted even if geoid is what was measured. The opposite is true for longwavelength phenomena.
Measurement of Marine Gravity Marine gravity measurements can be and have been acquired with several different sorts of sensors and from a variety of platforms, including ships, submarines, airplanes, and satellites. The ideal combination of sensor system and platform depends upon the needed accuracy, spatial coverage, and available time and funds. Gravimeters
The design for most marine gravimeters is borrowed from their terrestrial counterparts and are either absolute or relative in their measurements. Absolute gravimeters measure the full acceleration of gravity g at the survey site along the direction of the local vertical. Modern marine absolute gravimeters measure precisely the vertical position z of a falling mass (e.g., a corner cube reflector) as a function of time t in a vacuum cylinder using laser interferometry and an atomic clock. The acceleration is then calculated as the second derivative of the position of the falling mass as a function of time (eqn [3]). g¼
d2 zðtÞ dt2
½3
Absolute gravimeters tend to be larger, more difficult to deploy, costlier to build, and more expensive to run than relative gravimeters, and thus are only used when relative gravimeters are inadequate for the problem being addressed. Most gravity measurements at sea are relative measurements, Dg: the instrument measures the difference between gravity at the study site and at another site where absolute gravity is known (e.g., the dock where the expedition originates). Modern relative gravimeters are based on Hooke’s law for the force F required to extend a spring a distance x (eqn [4]), where ks is the spring constant, calculated by extending the spring under a known force. F ¼ ks x ¼ mg
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at other locations can be calculated by observing how much more or less that same spring is stretched at other locations by the same mass. Although such systems are relatively inexpensive to build and easy to deploy, they suffer from drift: in effect, the spring constant changes with time because no physical spring is perfectly elastic. To first order, the drift can be corrected by returning to the same or another base station with the same instrument, and assuming that the drift was linear with time in between. The accuracy of this linear drift assumption improves with more frequent visits to the base station, but this is usually impractical for marine surveys. Through clever design, the latest generation of marine gravimeters has greatly reduced the drift problem as compared with earlier instruments. The measurement of the gravity gradient tensor was widely used early in the twentieth century for oil exploration, but fell into disfavor in the 1930s as scalar gravimeters became more reliable and easy to use. Gravity gradiometry at sea is currently making a comeback as the result of declassification of military gradiometer technology developed for use in submarines during the Cold War. Gravity gradiometers measure the three-dimensional gradient in the gravity vector using six pairs of aligned gravimeters, with accuracies reaching better than 1 Eo¨tvo¨s. In comparison with measurements of gravity, the gravity gradient has more sensitivity to variations at short wavelenghts (B5 km or less), making it useful for delimiting shallow structures buried beneath the seafloor. Geoid anomalies can be directly measured from orbiting satellites carrying radar altimeters. The altimeters measure the travel time of a radar pulse from the satellite to the ocean surface, from which it is reflected and bounced back to the satellite. Tracking stations on Earth solve for the position of the satellite with respect to the center of the Earth. These two types of information are then combined to calculate the height of the sea level equipotential surface above the center of the Earth. Because the solid land surface does not follow an equipotential, altimeters cannot be used to constrain the terrestrial geoid. Furthermore, it is difficult to extract geoid from ocean areas covered by sea ice. However, in the near future, laser altimeters deployed from satellites hold the promise of extracting geoid information even over ocean surfaces marred with sea ice, on account of their enhanced resolution.
½4
If a mass m is suspended from this spring at a site where gravity g is known (e.g., by deploying an absolute gravimeter at that base station), then gravity
Platforms
Marine gravity data can be acquired either from moving or from stationary platforms. Because the
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gravity field from variations in the depth of the sea floor is such a large component of the observed signal, most marine gravity surveys have relied on ships or submarines that enable the simultaneous acquisition of depth observations. However, airborne gravity measurements have been acquired successfully over ice-covered areas of the polar oceans, and orbiting satellites have measured the marine geoid from space. A major challenge in acquiring gravity data from a floating platform at the sea surface is in separating the acceleration of the platform in the dynamic ocean from the acceleration of gravity. This problem is overcome by mounting marine gravimeters deployed from ships on inertially stabilized tables. These tables employ gyroscopes to maintain a constant attitude despite the pitching and rolling of the ship beneath the table. The nongravitational acceleration is somewhat mitigated by mounting marine gravimeters deep in the hold and as close to the ship’s center of motion as possible. Special damping mechanisms also prevent the spring in the gravimeter from responding to extremely high-frequency changes in the force on the suspended mass. Instruments deployed in submersibles resting on the bottom of the ocean or in instrument packages lowered to the bottom of the ocean do not suffer from the dynamic accelerations of the moving ocean surface, but bottom currents can also be an important source of noise in submarine gravimetry. Installing instruments in boreholes is the most effective way to counter this problem, but it is also an expensive solution.
Reduction of Marine Data A number of standard corrections must be applied to the raw gravity data (either g or Dg) prior to interpretation. In addition to any drift correction, as mentioned above for relative gravity measurements, a latitude correction is immediately applied to account for the large change in gravity between the poles and the Equator caused by Earth’s rotation. Near the Equator, the centrifugal acceleration from the Earth’s spin is large, and gravity is about 50 000 gu less, on average, than at the poles. Because this effect is 5000 times larger than typical regional gravity signals of interest, it must be removed from the data using a standard formula for the variation of gravity g0 on a spheroid of revolution best fitting the shape of the Earth (eqn [5]; Y ¼ latitude). g0 ðYÞ ¼ 9:7803185 1 þ 5:278895 103 sin2 Y þ2:3462 105 sin4 Y ms2 ½5
A second correction that must be made if the gravity is measured from a moving vehicle, such as a ship or airplane, accounts for the effect on gravity of the motion of the vehicle with respect to the Earth’s spin. A ship steaming to the east is, in effect, rotating faster than the Earth. The centrifugal effect of this increased rate of rotation causes gravity to be less than it would be if the ship were stationary. The opposite effect occurs for a ship steaming to the west. This term, called the Eo¨tvo¨s correction, is largest near the Equator and involves only the east–west component vEW of the ship’s velocity vector (eqn [6]), in which o is the angular velocity of the Earth’s rotation. gEOT ¼ 2ovEW cos Y
½6
The free air gravity correction, which accounts for the elevation of the measurement above the Earth’s sea level equipotential surface, is obviously not needed if the measurement is made on the sea surface. The free air correction gFA is required if the measurement is made from a submersible or an airplane: eqn [7], where h is elevation above sea level in meters. gFA ¼ 3:1h
gu
½7
This correction is added to the observation if the sensor is deployed above the Earth’s surface, and subtracted for stations below sea level. For land surveys, the Bouguer correction accounts for the extra mass of the topography between the observation and sea level. For its marine equivalent, it adds in the extra gravitational attraction that would be present if rock rather than water existed between sea level and the bottom of the ocean. Except in areas of rugged bathymetry, the Bouguer correction gB is calculated using the slab formula (eqn [8]). gB ¼ 2pDrGz
½8
Here Dr is the density difference between oceanic crust and sea water, G is Newton’s constant, and z is the depth of the sea floor. This correction is seldom used because it produces very large positive gravity anomalies. Furthermore, there are more accurate corrections for the effect of bathymetry that do not make the unrealistic assumption that the expected state for the oceans should be that the entire depth is filled with crustal rocks displacing the water. The Bouguer correction is necessary, however, when gravity measurements are made from a submarine, in order to combine those data with more conventional observations from the sea surface. In this case, the Bouguer correction is applied twice: once to remove
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GRAVITY
the upward attraction of the layer of water above the submarine, and once more to add in that layer’s gravitation field below the sensor. Satellite measurements of sea surface height go through a different processing sequence to recover marine geoid anomalies. The most important step is in calculating precise orbits. Information from tracking stations is supplemented with a ‘crossover analysis’ that removes long-wavelength bias in orbit elevation by forcing the height valves to agree wherever orbits cross. Corrections are then made for known physical oceanographic effects such as tides, and wave action is averaged out. The height of the sea level geoid above the Earth’s center, assuming the standard spheroid, is subtracted from the data to create geoid anomalies.
History A principal impediment to the acquisition of useful gravity observations at sea was the difficulty in separating the desired acceleration of gravity from the acceleration of the platform floating on the surface of the moving ocean. For this reason, the first successful gravity measurements to be acquired at sea were taken from a submarine by the Dutch pioneer, Vening Meinesz, in 1923. He used a pendulum gravimeter, which was the state of the art for measuring absolute gravity at that time. By accurately timing the period, T, of the swinging pendulum, the acceleration of gravity, g, can be recovered according to eqn [9], in which l is the length of the pendulum arm. sffiffiffi l T ¼ 2p g
½9
By 1959, five thousand gravity measurements had been acquired from submarines globally. These measurements were instrumental in revealing the large gravity anomalies associated with the great trenches along the western margin of the Pacific. However, these gravity observations were very timeconsuming to acquire because of the long integration times needed to achieve a high-precision estimate of the pendulum’s period, and could not be adapted for use on a surface ship. Gravity measurements at sea became routine and reliable in the late 1950s with the development of gyroscopically stabilized platforms and heavily damped mass-and-spring systems constrained to move only vertically. The new platforms compensated for the pitch and roll of the ship such that simple mass-and-spring gravimeters could collect
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time series of variations in gravity over the oceans from vessels under way. Without any need to stop the ship on station, a time series of gravity measurements could be obtained at only small incremental cost to ship operations. With the advent of the new instrumentation, the catalogue of marine gravity values has grown in the past 40 years to more than 2.5 million measurements. A new era of precision in marine gravity began with the advent of the Global Positioning System (GPS) in the late 1980s. Prior to this time, the largest source of uncertainty in marine gravity lay in the Eo¨tvo¨s correction. Older navigation systems (dead reckoning, celestial, and even the TRANSIT satellite system) were too imprecise in the absolute position of the ship and too infrequently available to allow accurate velocity estimation from minute to minute, especially if the ship was maneuvering. Typically, gravity data had to be discarded for an hour or so near the time of any change in course. The high positioning accuracy and frequency of GPS fixes now allows such precise calculation of the Eo¨tvo¨s correction that it is no longer the limiting factor in the accuracy of marine gravity data. A breakthough in determining the global marine gravity field was achieved with the launching of the GEOS-3 (1975–1977) and Seasat (1978) satellites, which carried radar altimeters. Altimeters were deployed for the purpose of measuring dynamic sea surface elevation associated with physical oceanographic effects. The Seasat satellite carried a new, high-precision altimeter that characterized the variations in sea surface elevation with unprecedented detail. The satellite failed prematurely, but not before it returned a wealth of data on the marine gravity field from its observations of the marine geoid. The geoid variations at mid- and short-wavelength were so large that the dynamic oceanographic effects motivating the mission could be considered a much smaller noise term. The success of the Seasat mission led to the launch of Geosat, which measured the geoid at even higher precision and resolution. Unfortunately, most of that data remained classified by the US military until the results from a similar European mission were about to be released into the public domain. The declassification of the Geosat data in 1995 fueled a major revolution in our understanding of the deep seafloor (Figure 1). The latest developments in marine gravity stem from the desire to detect the shortest spatial wavelengths of gravity variations by taking gravimeters to the bottom of the ocean. Gravity is one example of a potential field, and as such the amplitude, A, of the signal of interest decays with distance, z, between source and detector as in eqn [10], where k is the
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Figure 1 Gravity field over the North Pacific. This view is based on satellite altimetry data from the Geosat and other missions. (Data from Sandwell and Smith (1997).)
modulus of the spatial wavenumber, the reciprocal of the spatial wavelength. ABe2pkz
½10
For sensors located on a ship at the sea surface in average ocean depths of 4.5 km, it is extremely difficult to detect short-wavelength variations in gravity of a few kilometers or less. Even lowering the gravimeter to the cruising depth of most submarines (a few hundred meters) does little to overcome the upward attenuation of the signal from localized sources on and beneath the seafloor. The solution to this problem recently has been to take gravimeters to the bottom of the ocean, either in a deep-diving submersible such as Alvin, or as an instrument package lowered on a cable. Most gravity measurements at sea are relative measurements. However, recent advances in instrumentation now allow absolute gravimeters to be deployed on the bottom of the ocean, avoiding the problem of instrument drift that adds error to relative gravity measurements. However, noise associated with the short baselines required for operation in the deep sea remains problematic.
Interpretation of Marine Gravity Short-Wavelength Anomalies
The shortest-wavelength gravity anomalies over the oceans (less than a few tens of kilometers) are the
least ambiguous to interpret since they invariably are of shallow origin. The upward continuation factor guarantees that any spatially localized anomalies with deep sources will be undetectable at the ocean surface. Near-bottom gravity measurements are able to improve somewhat the detection of concentrated density anomalies buried at deeper levels, but most are assumed to lie within the oceanic crust. One of the most useful applications of shortwavelength gravity anomalies has been to predict ocean bathymetry (Figure 2). Radar altimeters deployed on the Seasat and Geosat missions measured with centimeter accuracy the height of the underlying sea surface, an excellent approximation to the marine geoid, over all ice-free marine regions. The accuracy and spatial coverage was far better than had been provided from more than a century of marine surveys from ships. At short wavelengths, undulations of the rock–water interface are the largest contribution to the short-wavelength portion of the geoid spectrum, which opened up the possibility of predicting ocean depth from the excellent geoid data. For example, an undersea volcano, or seamount, represents a mass excess over the water it displaces. The extra mass locally raises the equipotential surface, such that positive geoid anomalies are seen over volcanoes and ridges while geoid lows are seen over narrow deeps and trenches. The prediction of bathymetry from marine geoid or gravity data is tricky: the highest frequencies in the bathymetry cannot be estimated because of the upward attenuation
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GRAVITY
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(A)
(B)
Figure 2 Example of bathymetric prediction from gravity anomalies in a largely unexplored region of the South Pacific. (A) The best available bathymetry from sparse echo soundings available in the early 1990s. (B) A diagram shows a dramatic improvement in definition of the bathymetry when satellite gravity observations are used to constrain the short-wavelength component of the bathymetry. (Adapted from McNutt and Bonneville (1996).)
problem, and the longer wavelengths are canceled out in the geoid by their isostatic compensation (see following section). These longer wavelengths in the bathymetry must be introduced into the solution using traditional echo soundings from sparse ship tracks. Nevertheless, the best map we currently have of the depth of the global ocean is courtesy of satellite altimetry. Mid-Wavelength Anomalies
The mid-wavelength part of the gravity spectrum (tens to hundreds of kilometers) is dominated by the effects of isostatic compensation. Isostasy is the process by which the Earth supports variations in topography or bathymetry in order to bring about a condition of hydrostatic equilibrium at depth. The definition of isostasy can be extended to include both static and dynamic compensation mechanisms, but at these wavelengths the static mechanisms are most important. There are a number of different types of isostatic compensation at work in the oceans, and the details of the gravity field can be used to distinguish them and to estimate the thermomechanical behavior of oceanic plates. One of the simplest mechanisms for isostatic compensation is Airy isostasy: the oceanic crust is thickened beneath areas of shallow bathymetry. The thick crustal roots displace denser mantle material, such that the elevated features float on the mantle much like icebergs float in the ocean. Of the various
methods of isostatic compensation, this mechanism predicts the smallest gravity anomalies over a given feature. From analysis of marine gravity, we now know that this sort of compensation mechanism is only found where the oceanic crust is extremely weak, such as on very young lithosphere near a midocean ridge. For example, large plateaus formed when hotspots intersect midocean ridges are largely supported by Airy-type isostasy. Elsewhere the oceanic lithosphere is strong enough to exhibit some lateral strength in supporting superimposed volcanoes and other surface loads. An extremely common form for support of bathymetric features in the oceans is elastic flexure. Oceanic lithosphere has sufficient strength to bend elastically, thus distributing the weight of a topographic feature over an area broader than that of the feature itself (Figure 3). Analysis of marine gravity has been instrumental in establishing that the elastic strength of the oceanic lithosphere increases with increasing age. Young lithosphere near the midocean ridge is quite weak, in some cases hardly distinguished from Airytype isostasy. The oldest oceanic lithosphere displays an effective thickness equivalent to that of a perfectly elastic plate 40 km thick. The fact that this thickness is less than that of the commonly accepted value for the thickness of the mechanical plate that drifts over the asthenosphere indicates that the base of the oceanic lithosphere is not capable of sustaining large deviatoric stresses (of the order of 100 MPa or more) over million-year timescales.
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Net gravity anomaly Contribution from volcano
Contribution from flexure of plate Volcano
Light crust
Elastic plate Dense mantle
Figure 3 Cartoon showing how the seafloor is warped as an elastic plate under the weight of a small volcano. The gravity anomaly that would be detected by a ship sailing along the sea surface over this feature is the net difference between the positive gravity perturbation from the extra mass of the volcano and the negative gravity perturbation produced when the elastically flexed light crust replaces denser mantle. (Adapted from McNutt and Bonneville (1996).)
Another very important method of isostatic compensation in the ocean is Pratt isostasy. This method of support supposes that the height of a vertical column of bathymetry is inversely proportional to its density. Low-density columns can be higher because they are lighter, whereas heavy columns must be short in order to produce the same integrated mass at some assumed depth of compensation. In the oceans, variations in the temperature of the lithosphere produce elevation changes in the manner of Pratt isostasy. For example, ridges stand 4 km above the deep ocean basins because the underlying lithosphere is hotter when the plate is young. The bathymetric swells around young hotspot volcanoes may also be supported by Pratt-type isostasy, although some combination of crustal thickening and dynamic isostasy may be operating as well. Again, gravity and geoid anomalies have been principal constraints in arguing for the mechanism of support for bathymetric swells. Long-Wavelength Anomalies
At wavelengths from 1000 to several thousand kilometers, gravity anomalies are usually derived from satellite observations and interpreted using equations appropriate for a spherical earth. Geoid is interpreted more commonly than gravity directly, as it emphasizes the longer wavelengths in the geopotential field. Isostatic compensation for smaller-scale bathymetric features, such as seamounts, can usually be ignored in that the gravity anomaly from bathymetry is canceled out by that from its compensation when spatially averaged over longer wavelengths.
The principal signal at these wavelengths arises from the subduction of lithospheric slabs and other sorts of convective overturn within the mantle. Three sorts of gravitational contributions must be considered: (1) the direct effect of mass anomalies within the mantle, either buoyant risers or dense sinkers which drive convection; (2) the warping of the surface caused by viscous coupling of the risers or sinkers to the earth’s surface; and (3) the warping of any deeper density discontinuities (such as the core– mantle boundary) also caused by viscous coupling. In the 1980s, estimates of the locations and densities of mass anomalies in the mantle responsible for the first contribution above began to become available courtesy of seismic tomography. Travel times of earthquake waves constrained the locations of seismically fast and slow regions in the mantle. By assuming that the seismic velocity variations were caused by temperature differences between hot, rising material and cold, sinking material, it was possible to convert velocity to density using standard relations. Knowing the locations of the mass anomalies driving convection inside the Earth led to a breakthrough in understanding the long-wavelength gravity and geoid fields. The amount of deformation on density interfaces above and below the mass anomalies inferred from tomography (contributions (2) and (3) above) depends upon the viscosity structure of Earth’s mantle. Coupling is more efficient with a more viscous mantle, whereas a weaker mantle is able to soften the transmission of the viscous stresses from the risers and sinkers. Therefore, one of the principal uses of marine gravity anomalies at long wavelengths has been to calibrate the viscosity structure of the oceanic upper mantle. This interpretation must be constrained by estimates of the dynamic surface topography over the oceans, which is actually easier to estimate than over the continents because of the relatively uniform thickness of oceanic crust. A fairly common result from this sort of analysis is that the oceanic upper mantle must be relatively inviscid. The geoid shows that there are large mass anomalies within the mantle driving convection that are poorly coupled to variations in the depth of the seafloor. If the upper mantle were more viscous, there should be a stronger positive correlation between marine geoid and depth of the seafloor at long wavelengths.
See also Manned Submersibles, Deep Water. Satellite Altimetry. Satellite Oceanography, History and Introductory Concepts.
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Further Reading Bell RE, Anderson R, and Pratson L (1997) Gravity gradiometry resurfaces. The Leading Edge 16: 55--59. Garland GD (1965) The Earth’s Shape and Gravity. New York: Pergamon Press. McNutt MK and Bonneville A (1996) Mapping the seafloor from space. Endeavour 20: 157--161. McNutt MK and Bonneville A (2000) Chemical origin for the Marquesas swell. Geochemistry, Geophysics and Geosystems 1. Sandwell DT and Smith WHF (1997) Marine gravity for Geosat and ERS-1 altimetry. Journal of Geophysical Research 102: 10039--10054. Smith WHF and Sandwell DT (1997) Global seafloor topography from satellite altimetry and ship depth soundings. Science 277: 1956--1962.
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Turcotte DL and Schubert G (1982) Geodynamics: Applications of Continuum Physics to Geological Problems. New York: Wiley. Watts AB, Bodine JH, and Ribe NM (1980) Observations of flexure and the geological evolution of the Pacific Ocean basin. Nature 283: 532--537. Wessel P and Watts AB (1988) On the accuracy of marine gravity measurements. Journal of Geophysical Research 93: 393--413. Zumberg MA, Hildebrand JA, Stevenson JM, et al. (1991) Submarine measurement of the Newtonian gravitational constant. Physical Review Letters 67: 3051--3054. Zumberg MA, Ridgeway JR, and Hildebrand JA (1997) A towed marine gravity meter for near-bottom surveys. Geophysics 62: 1386--1393.
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GROUNDWATER FLOW TO THE COASTAL OCEAN A. E. Mulligan and M. A. Charette, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Water flowing through the terrestrial landscape ultimately delivers fresh water and dissolved solutes to the coastal ocean. Because surface water inputs (e.g., rivers and streams) are easily seen and are typically large point sources to the coast, they have been well studied and their contributions to ocean geochemical budgets are fairly well known. Similarly, the hydrodynamics and geochemical importance of surface estuaries are well known. Only recently has significant attention turned toward the role of groundwater inputs to the ocean. Historically, such inputs were considered insignificant because groundwater flow is so much slower than riverine flow. Recent work however has shown that groundwater flow through coastal sediments and subsequent discharge to the coastal ocean can have a significant impact on geochemical cycling and it is therefore a process that must be better understood. Groundwater discharge into the coastal ocean generally occurs as a slow diffuse flow but can be found as large point sources in certain terrain, such as karst. In addition to typically low flow rates, groundwater discharge is temporally and spatially variable, complicating efforts to characterize site-specific flow regimes. Nonetheless, the importance of submarine groundwater discharge (SGD) as a source of dissolved solids to coastal waters has become increasingly recognized, with recent studies suggesting that SGDderived chemical loading may rival surface water inputs in many coastal areas. So while the volume of water discharged as SGD may be small relative to surface discharge, the input of dissolved solids from SGD can surpass that of surface water inputs. For example, SGD often represents a major source of nutrients in estuaries and embayments. Excess nitrogen loading can result in eutrophication and its associated secondary effects including decreased oxygen content, fish kills, and shifts in the dominant flora. First, let us define ‘groundwater’ in a coastal context. We use the term to refer to any water that resides in the pore spaces of sediments at the land–ocean boundary. Hence, such water can be fresh terrestrially derived water that originates as
88
rainwater infiltrating through the subsurface or it can represent saline oceanic water that flows through the sediments (Figure 1). Therefore, groundwater discharging to coastal waters can have salinity that spans a large range, being some mixture of the two end members. We therefore use the terms fresh SGD and saline SGD to distinguish these sources of fluid and brackish SGD to mean a mixture of the fresh and saline end members.
Basics of Groundwater Flow Groundwater flow in the subsurface is driven by differences in energy – water flows from high energy areas to low energy. The mechanical energy of a unit volume of water is determined by the sum of gravitational potential energy, pressure energy, and kinetic energy: Energy per unit volume ¼ rgz þ P þ
rV 2 2
½1
where r is fluid density, g is gravitational acceleration, z is elevation of the measuring point relative to a datum, P is fluid pressure at the measurement point, and V is fluid velocity. Because groundwater flows very slowly (on the order of 1 m day 1 or less), its kinetic energy is very small relative to its gravitational potential and pressure energies and the kinetic energy term is therefore ignored. By removing the kinetic energy term and rearranging eqn [1] to express energy in terms of mechanical energy per unit weight, the concept of hydraulic head is developed: Energy per unit weight ¼ hydraulic head ¼ z þ
P rg ½2
Groundwater therefore flows from regions of high hydraulic head to areas of low hydraulic head. Because groundwater flows through a porous media, the rate of flow depends on soil properties such as the degree to which pore spaces are interconnected. The property of interest in groundwater flow is the permeability, k, which is a measure of the ease with which a fluid flows through the soil matrix. Groundwater flow rate can then be calculated using Darcy’s law, which says that the flow rate is linearly proportional to the hydraulic gradient:
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q¼
rgk ðrhÞ m
½3
GROUNDWATER FLOW TO THE COASTAL OCEAN
89
Precipitation Evaporation
Overland flow Infiltratio
n
Septic system
Ocean
Seawater circulation
ate
r
Water table
tw
Groundwater
l sa r− ce e t a fa shw Inter Fre
Figure 1 Simplified water cycle at the coastal margin. Modified from Heath R (1998) Basic ground-water hydrology. US Geological Survey Water Supply Paper 2220. Washington, DC: USGS.
where q is the Darcy flux, or flow rate per unit surface area, and m is fluid viscosity. A more general expression of Darcy’s law is: k q ¼ ðrP þ rgrzÞ m
Runup and waves
z
½4
In inland aquifers, the density of groundwater is constant and eqn [4] is reduced to the simpler form of Darcy’s law (eqn [3]). In coastal aquifers, however, the presence of saline water along the coast means that the assumption of constant density is not valid and so the more inclusive form of Darcy’s law, eqn [4], is required.
Groundwater Flow at the Coast Several forces drive groundwater flow through coastal aquifers, leading to a complex flow regime with significant variability in space and time (Figure 2). The primary driving force of fresh SGD is the hydraulic gradient from the upland region of a watershed to the surface water discharge location at the coast. Freshwater flux is also influenced by several other forces at the coastal boundary that also drive seawater through the sediments. For example, seawater circulates through a coastal aquifer under the force of gravity, from oceanic forces such as
Fresh Tidal pumping 40z
Saline
Dispersive circulation
Figure 2 Simplification of an unconfined coastal groundwater system. Water flow is driven by the inland hydraulic gradient, tides, beach runup and waves, and dispersive circulation. Other forcing mechanisms can drive fluid through coastal sediments, including seasonal changes in recharge to the inland groundwater system and tidal differences across islands.
waves and tides, as a result of dispersive circulation along the freshwater–saltwater boundary within the aquifer, and from changes in upland recharge. Several other forcing mechanisms exist, but they are generally only present in specific settings. For example, tidal height differences across many islands can drive flow through the subsurface. All of these forcing mechanisms affect the rate of fluid flow for
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GROUNDWATER FLOW TO THE COASTAL OCEAN
both fresh and saline groundwater and are ultimately important in controlling the submarine discharge of both fluids. The analysis of coastal groundwater flow must account for the presence of both fresh- and saline water components. When appropriate, such as in regional-scale analysis or for coarse estimation purposes, an assumption can be made that there is a sharp transition, or interface, between the fresh and saline groundwater. While this assumption is not strictly true, it is often appropriate, and invoking it results in simplifying the analysis. For example, we can estimate the position of the freshwater–saltwater interface by assuming a sharp interface, no flow within the saltwater region, and only horizontal flow within the fresh groundwater. Invoking these assumptions means that the pressures at adjacent points along the interface on both the freshwater and saltwater sides are equal. Equating these pressures and rearranging, the depth to the interface can be calculated as follows:
Interface depth ¼
r1 z ¼ 40z r2 r1
½5
where r1 is the density of fresh water (1000 kg m 3) and r2 is the density of seawater (1025 kg m 3). This equation states that the depth of the interface is 40 times the elevation of the water table relative to mean sea level. While eqn [5] is only an approximation of the interface location, it is very helpful in thinking about freshwater and saltwater movement in response to changes in fresh groundwater levels. As recharge to an aquifer increases, water levels increase and the interface is pushed downward. This is also equivalent to pushing the interface seaward and the net effect is to force saltwater out of the subsurface and to replace it with fresh water. The opposite flows occur during times of little to no recharge or if groundwater pumping becomes excessive. While the sharp-interface approach is useful for conceptualizing flow at the coast, particularly in large-scale problems, the reality is more complex. Not only does the saline groundwater flow but also a zone of intermediate salinity extends between the fresh and saline end members, establishing what many refer to as a ‘subterranean estuary’. Like their surface water counterparts, these zones are hotbeds of chemical reactions. Because the water in the interface zone ultimately discharges into coastal waters, the flow and chemical dynamics within the zone are critically important to understand. Research into these issues has only just begun.
Detecting and Quantifying Submarine Groundwater Discharge As a first step in quantifying chemical loads to coastal waters, the amount of water flowing out of the subsurface must be determined. This is particularly challenging because groundwater flow is spatially and temporally variable. A number of qualitative and quantitative techniques have been developed to sample SGD, with each method sampling a particular spatial and temporal scale. Because of limitations with each sampling method, several techniques should be used at any particular site. Physical Approaches
Infrared thermography Infrared imaging has been used to identify the location and spatial variability of SGD by exploiting the temperature difference between surface water and groundwater at certain times of the year. While this technique is quite useful for identifying spatial discharge patterns, it has not yet been applied to estimating flow rates. An example of thermal infrared imagery is shown in Figure 3, an image of the head of Waquoit Bay, a small semi-enclosed estuary on Cape Cod, Massachusetts, USA. In late summer, the groundwater temperature is approximately 13 1C, whereas surface water is about 7–10 1C warmer. Locations of SGD can be seen in the infrared image as locations along the beach face with cooler temperature than the surrounding surface water. The image clearly shows spatial variability in SGD along the beach face, information that is extremely valuable in designing an appropriate field sampling campaign. Hydrologic approaches There are two hydrologic approaches to estimating SGD: the mass balance method and Darcy’s law calculation. Both methods are typically applied to estimating fresh groundwater discharge, although Darcy’s law can also be used to estimate saline flow into and out of the seafloor. To apply Darcy’s law, one must measure the soil permeability and hydraulic head at several locations (at least two) at the field site. Data must also be gathered to determine the cross-sectional flow area. The field data are then used with Darcy’s law to calculate a groundwater flow rate into the coastal ocean. The main disadvantages of this approach include the fact that permeability is highly heterogeneous, often ranging over several orders of magnitude, and so an ‘average’ value to use with Darcy’s law is seldom, if ever, well known. Furthermore, hydraulic head measurements require invasive, typically expensive, well installations. Finally, hydraulic head
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GROUNDWATER FLOW TO THE COASTAL OCEAN
91
Temp. (°C) 2.3 2.4−7.4 7.5−8 8.1−9 9.1−9.5 9.6−10 top o con graph tou ic r
10.1−10.5 10.6−11 11.1−11.5
3-m
11.6−12 12.1−12.5 12.6−13 13.1−13.5 13.6−14 14.1−14.5
15
12
18
9
9
11 11
9
11
18
14.6−15 14
15.1−15.5 15.6−16
25
Waquoit Bay
16.1−16.5
Beach/water line
16.6−17 17.1−17.5 0
50
100
200 m
17.6−18 18.1−18.5 18.6−19 19.1−19.5 19.6−20 20.1−20.5
Figure 3 Thermal infrared image of the head of Waquoit Bay, Massachusetts. The beach/water line is shown as a red curve. Light grays imply higher temperatures and darker shades show lower temperatures. Lower temperatures within the bay indicate regions influenced by SGD. The bars show average groundwater seepage rates as measured by manual seep meters from high tide to low tide. Numbers below the bars are average rates in cm d 1. Modified from Mulligan AE and Charette MA (2006) Intercomparison of submarine ground water discharge estimates from a sandy unconfined aquifer. Journal of Hydrology 327: 411–425.
is a point measurement and capturing the spatial variability therefore requires installing many wells. The primary advantage of this approach is that it is well established and easy to implement: head measurements are easy to collect once wells are installed and the flux calculations are simple. The mass balance approach to estimating SGD requires ascertaining all inputs and outputs of water flow, except SGD, through the groundwatershed. Assuming a steady-state condition over a specified time frame, the groundwater discharge rate is calculated as the difference between all inputs and all outputs. Implementing this approach can be quite simple or can result in complex field campaigns, but the quality of the data obviously affect the level of uncertainty. Even with extensive field sampling, water budgets are seldom known with certainty and
so should be used with that in mind. Furthermore, if the spatial and temporal variability of SGD is needed for a particular study, the mass balance approach is not appropriate. Direct measurements: Seepage meters SGD can be measured directly with seepage meters. Manual seepage meters (Figure 4) are constructed using the tops of 55 gallon drums, where one end is open and placed into the sediment. The top of the seepage meter has a valve through which water can flow; a plastic bag prefilled with a known volume of water is attached to the valve, so that inflow to or outflow from the sediments can be determined. After a set length of time, the bag is removed and the volume of water in the bag is measured. The change in water volume over the sampling period is
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GROUNDWATER FLOW TO THE COASTAL OCEAN
then used to determine the average flow rate of fluid across the water–sediment boundary over the length of the sampling period. These meters are very simple to operate, but they are manually intensive and are sensitive to wave disturbance and currents. Furthermore, they only sample a small flow area and so many meters are needed to characterize the spatial variability seen at most sites (Figure 3). Recently, several other technologies have been applied toward developing automated seep meters. Technologies include the heat-pulse method, continuous
Water surface
Chemical Tracers
Collection bag
Connectors
heat, ultrasonic, and dye dilution. These meters can be left in place for days and often weeks and will measure seepage without the manual intervention needed using traditional seepage meters. The trade-off with these meters is that they are expensive and therefore only a limited number are typically employed at any given time. An example of seepage measurements made using the dye dilution meter is shown in Figure 5. Note that the seepage rates are inversely proportional to tidal height. As the tide rises, the hydraulic gradient from land to sea is reduced, seepage slows, and the flow reverses, indicating that seawater is flowing into the aquifer at this location during high tide. Conversely, at low tide, the hydraulic gradient is at its steepest and SGD increases.
Seepage meter
Fluid flow Sediment
Figure 4 Graphic of a manual seepage meter deployed at the sediment–water interface. These meters can be constructed using the top of a 55 gallon drum. The collection bag serves as a fluid reservoir so that both inflow to and outflow from the sediments can be measured. Modified from Lee DR (1977) A device for measuring seepage in lakes and estuaries. Limnology and Oceanography 22: 140–147.
The chemical tracer approach to quantifying SGD has an advantage over seepage meters in that it provides an integrated flux over a wide range of spatial scales from estuaries to continental shelves. The principle of using a chemical tracer is simple: find an element or isotope that is highly enriched (or depleted) in groundwater relative to other sources of water, like rivers or rainfall, to the system under study. If SGD is occurring, then the flux of this element via groundwater will lead to enrichment in the coastal zone that is well above background levels in the open ocean (Figure 6). A simple mass balance/box model for the system under study can be performed, where all sources of the tracer other than groundwater are subtracted from the total inventory of the chemical. The residual inventory, or ‘excess’, is then
75
Flow (cm d−1) Tide (cm)
50 Submarine groundwater discharge
25
0 −25 −50 6 Sep.
Flow Tide 7 Sep.
Seawater flow into sediments
8 Sep.
Date 2002 Figure 5 Seepage rates at Waquoit Bay, Massachusetts, USA, measured using a dye dilution automated seepage meter. Modified from Sholkovitz ER, Herbold C, and Charette MA (2003) An automated dye-dilution based seepage meter for the time-series measurement of submarine groundwater discharge. Limnology and Oceanography: Methods 1: 17–29.
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GROUNDWATER FLOW TO THE COASTAL OCEAN
34 Myrtle 20 Beach 23 20 17 15 Inner 13 shelf 12 10 14 15
(dpm 100 l−1) Jul. 1994
226Ra
Winyah Bay
17
25 20 15 Inner 13 shelf 11 10 14 13
Sh
am
re
t lf S
Gu
Open ocean = 8.0 ± 0.5
am
tre
Sh
12 11 10
31 81
ak
re
b elf
8.9
el
32
11 8.9
ak
20 18
15 15 11 9.3 8.7 10 10
re
26
17
Cape Fear
10
17 19 16 13 20 19 10 25 8.2 Charleston 28 24 21 7.8 10 24 20 8.7 20 11 14 11 25 15 10 10 9.7 19 11 11 11 10 9.4 10 11
fb
Sav an Riv nah er
Latitude (°N)
33
25
93
S ulf
0
20
40
G
80
79
78
60 km
80
100
77
Longitude (° W) Figure 6 The distribution of 226Ra offshore South Carolina reveals high activities on the inner shelf that decrease offshore. Moore (1996) used the excess 226Ra to estimate regional SGD fluxes.
divided by the concentration of the tracer in the discharging groundwater to calculate the groundwater flow rate. Naturally occurring radionuclides such as radium isotopes and radon-222 have gained popularity as tracers of SGD due to their enrichment in groundwater relative to other sources and their built-in radioactive ‘clocks’. The enrichment of these tracers is owed to the fact that the water–sediment ratio in aquifers is usually quite small and that aquifer sediments (and sediments in general) are enriched in many U and Th series isotopes; while many of these isotopes are particle reactive and remain bound to the sediments, some like Ra can easily partition into the aqueous phase. Radon-222 (t1/2 ¼ 3.82 days) is the daughter product of 226Ra (t1/2 ¼ 1600 years) and a noble gas; therefore, it is even more enriched in groundwater than radium.
A key issue when comparing different techniques for measuring SGD is the need to define the fluid composition that each method is measuring (i.e., fresh, saline, or brackish SGD). For example, whereas hydrogeological techniques are estimates of fresh SGD, the radium and radon methods include a component of recirculated seawater. Therefore, it is often not possible to directly compare the utility of these techniques. Instead, they should be regarded as complementary. One of the seminal studies that showed how SGD can impact chemical budgets on the scale of an entire coastline was conducted off North and South Carolina, USA. Literature estimates of water residence time, riverine discharge/suspended sediment load, and the activity of desorbable 226Ra on riverine particles were used to determined that only B50% of the 226Ra inventory on the inner continental shelf
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GROUNDWATER FLOW TO THE COASTAL OCEAN
222Rn
140
7
120
6
100
5
80
4
60
3
40
2
20
1 18 May
19 May
20 May
40
30
0 21 May
20
10
Seep meter (cm d−1)
8
Water level (cm)
(dpm l−1)
94
0
Date 2002 Tide Seep meter Radon Figure 7 Plot comparing variations in SGD at Shelter Island, NY, between an automated seepage meter and the continuous radon method.
off North and South Carolina could be explained by surface inputs (Figure 6). The remaining inventory enters the system via SGD. Using an estimate of groundwater 226Ra, it is calculated that the groundwater flux to this region of the coastline is on the order of 40% of the river water flux. The approach for quantifying SGD using 222Rn is similar to that for radium (226Ra), except for a few key differences: (1) 222Rn loss to the atmosphere must be accounted for in many situations, (2) there is no significant source from particles in rivers, and (3) decay must be accounted for owing to its relatively short half-life. The first example of 222Rn use to quantify SGD to the coastal zone occurred in a study in the northeastern Gulf of Mexico. The strong pycnocline that develops in the summer time means that fluid that flows from the sediments into the bottom boundary layer is isolated from the atmosphere and therefore no correction for the air–sea loss of 222Rn is needed. Using this approach, diffuse SGD in a 620 km2 area of the inner shelf was estimated to be equivalent to B20 first-magnitude springs (a firstmagnitude spring has a flow equal to or greater than 245 000 m3 water per day, which is equivalent to B60% of the daily water supply for the entire state of Rhode Island, USA). Since 2000, a number of SGD estimation technique intercomparison experiments have been conducted through a project sponsored by the Scientific Committee on Oceanic Research (SCOR) and the Land– Ocean Interaction in the Coastal Zone (LOICZ)
Project. During these intercomparisons, several methods (chemical tracers, different types of seep meters, hydrogeologic approaches, etc.) were run side by side to evaluate their relative strengths and weaknesses. Figure 7 displays a comparison from the Shelter Island, NY, experiment of calculated radon fluxes (based on measurements from a continuous radon monitor), with seepage rates measured directly via the dye dilution seepage meter. During the period (17–20 May) when both devices were operating at the same time, there is a clear and reproducible pattern of higher radon and water fluxes during the low tides. There is also a suggestion that the seepage spikes slightly led the radon fluxes, which is consistent with the notion that the groundwater seepage is the source of the radon (the radon monitor was located offshore of the seepage meter). The excellent agreement in patterns and overlapping calculated advection rates (seepage meter: 2–37 cm day 1, average ¼ 127 8 cm day 1; radon model: 0–34 cm day 1, average ¼ 1177 cm day 1) by these two completely independent assessment tools is reassuring.
Geochemistry of the Subterranean Estuary The magnitude of chemical fluxes carried by SGD is influenced by biogeochemical processes occurring in the subterranean estuary, defined as the mixing zone between groundwater and seawater in a coastal
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GROUNDWATER FLOW TO THE COASTAL OCEAN
95
(a)
Fe (ppm) 4500
0
0
1500
Fe (ppm)
3000
Tan/red Tan/gray
100
150
Dark red
150
50
0
50
100 P (ppm)
150
1500
3000
50
100
Core 3 200
4500
150
Core 2 200
0
0
Yellow
Depth (cm)
Orange Red/orange
100
50
4500
Gray
3000
Gray
1500
Gray
0
0
Gray/yellow
Fe (ppm) (b)
200
Core 5 200
0
50
100
150
200
P (ppm)
0
50
100
150
200
P (ppm)
Figure 8 (a) Scientists extruding a sediment core taken through the subterranean estuary of Waquoit Bay, MA. Note the presence of iron oxides within the sediments at the bottom of the core (orange-stained sediments in foreground). (b) Changes in iron and phosphate concentration with depth in three sediment cores similar to the one shown in (a). The red circles indicate Fe concentration (ppm:mg Fe/g dry sediment) while the blue diamonds represent P (ppm:mg P/g dry sediment). Error bars indicate the standard deviation for triplicate leaches performed on a selected number of samples. The dashed lines represent the concentration of Fe and P in ‘off-site’ quartz sand. Also shown is the approximate color stratigraphy for each core. The R2 value for Fe vs. P in cores 2, 3, and 5 is 0.80, 0.91, and 0.16, respectively.
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GROUNDWATER FLOW TO THE COASTAL OCEAN
aquifer. The biogeochemical reactions in such underground estuaries are presumed to be similar to the surface estuary (river water/seawater) counterpart, though few comprehensive studies of chemical cycling in subterranean estuaries have been undertaken. Drivers of biogeochemical reactions in these environments include oxidation–reduction gradients, desorption–sorption processes, and microbially driven diagenesis. In a study of the Waquoit Bay subterranean estuary, a large accumulation of iron (hydr)oxide-coated sediments within the fresh–saline interface was encountered. These iron-oxide-rich sands could act as a geochemical barrier by retaining and accumulating certain dissolved chemical species carried to the subterranean estuary by groundwater and/or coastal seawater. Significant accumulation of phosphorus in the iron oxide zones of the Waquoit cores exemplifies this process (Figure 8). Phosphorous is not the only nutrient that can be retained/removed via reactions in the subterranean estuary. The microbial reduction of nitrate to inert dinitrogen gas, a process known as denitrification, is known to occur in the redox gradients associated with fresh and saline groundwater mixing. Conversely, ammonium, which is more soluble in saline environments, may be released within the subterranean estuary’s mixing zone. While the overall importance of SGD on the ‘global cycle’ of certain chemical species remains to be seen, there is little doubt that SGD is important at the local scale both within the United States and throughout the world.
Summary Groundwater discharge to the coastal ocean can be an important source of fresh water and dissolved solutes. Although a significant amount of research into the role of SGD in solute budgets has only occurred in the past decade or so, there is increasing evidence that solute loading from groundwater can be significant enough to affect solute budgets and even ecosystem health. Proper estimation of solute loads from groundwater requires confident estimates of both the groundwater discharge rate and average solute concentrations in the discharging fluid, neither of which are easily determined. Estimating groundwater discharge rates is complicated by the spatial and temporal variability of groundwater flow. A multitude of time-varying driving mechanisms complicate analysis, as does geologic heterogeneity. Nonetheless, a suite of tools from hydrogeology, geophysics, and geochemistry have been developed for sampling and measuring SGD.
The nature of geochemical reactions within nearshore sediments is not well understood, yet recent studies have shown that important transformations occur over small spatial scales. This is an exciting and critically important area of research that will reveal important process in the near future.
See also Chemical Processes in Estuarine Sediments. Estuarine Circulation. Gas Exchange in Estuaries. Inverse Modeling of Tracers and Nutrients. LongTerm Tracer Changes. Submarine Groundwater Discharge. Trace Element Nutrients. Tracer Release Experiments. Tracers of Ocean Productivity.
Further Reading Burnett WC, Aggarwal PK, Bokuniewicz H, et al. (2006) Quantifying submarine groundwater discharge in the coastal zone via multiple methods. Science of the Total Environment 367: 498--543. Burnett WC, Bokuniewicz H, Huettel M, Moore WS, and Taniguchi M (2003) Groundwater and pore water inputs to the coastal zone. Biogeochemistry 66: 3--33. Cable JE, Burnett WC, Chanton JP, and Weatherly GL (1996) Estimating groundwater discharge into the northeastern Gulf of Mexico using radon-222. Earth and Planetary Science Letters 144: 591--604. Charette MA and Sholkovitz ER (2002) Oxidative precipitation of groundwater-derived ferrous iron in the subterranean estuary of a coastal bay. Geophysical Research Letters 29: 1444. Charette MA and Sholkovitz ER (2006) Trace element cycling in a subterranean estuary. Part 2: Geochemistry of the pore water. Geochimica et Cosmochimica Acta 70: 811--826. Heath R (1998) Basic ground-water hydrology. US Geological Survey Water Supply Paper 2220. Washington, DC: USGS. Kohout F (1960) Cyclic flow of salt water in the Biscayne aquifer of southeastern Florida. Journal of Geophysical Research 65: 2133--2141. Lee DR (1977) A device for measuring seepage in lakes and estuaries. Limnology and Oceanography 22: 140--147. Li L, Barry DA, Stagnitti F, and Parlange J-Y (1999) Submarine groundwater discharge and associated chemical input to a coastal sea. Water Resources Research 35: 3253--3259. Michael HA, Mulligan AE, and Harvey CF (2005) Seasonal water exchange between aquifers and the coastal ocean. Nature 436: 1145--1148 (doi:10.1038/ nature03935). Miller DC and Ullman WJ (2004) Ecological consequences of ground water discharge to Delaware Bay, United States. Ground Water 42: 959--970.
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GROUNDWATER FLOW TO THE COASTAL OCEAN
Moore WS (1996) Large groundwater inputs to coastal waters revealed by Ra-226 enrichments. Nature 380: 612--614. Moore WS (1999) The subterranean estuary: A reaction zone of ground water and sea water. Marine Chemistry 65: 111--125. Mulligan AE and Charette MA (2006) Intercomparison of submarine ground water discharge estimates from a sandy unconfined aquifer. Journal of Hydrology 327: 411--425. Paulsen RJ, Smith CF, O’Rourke D, and Wong T (2001) Development and evaluation of an ultrasonic ground water seepage meter. Ground Water 39: 904--911. Portnoy JW, Nowicki BL, Roman CT, and Urish DW (1998) The discharge of nitrate-contaminated groundwater from developed shoreline to marsh-fringed estuary. Water Resources Research 34: 3095--3104.
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Shinn EA, Reich CD, and Hickey TD (2002) Seepage meters and Bernoulli’s revenge. Estuaries 25: 126--132. Sholkovitz ER, Herbold C, and Charette MA (2003) An automated dye-dilution based seepage meter for the timeseries measurement of submarine groundwater discharge. Limnology and Oceanography: Methods 1: 17--29. Slomp CP and Van Cappellen P (2004) Nutrient inputs to the coastal ocean through submarine groundwater discharge: Controls and potential impact. Journal of Hydrology 295: 64--86. Taniguchi M, Burnett WC, Cable JE, and Turner JV (2002) Investigation of submarine groundwater discharge. Hydrological Processes 16: 2115--2129. Valiela I, Foreman K, LaMontagne M, et al. (1992) Couplings of watersheds and coastal waters: Sources and consequences of nutrient enrichment in Waquoit Bay, Massachusetts. Estuaries 15: 443--457.
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HABITAT MODIFICATION M. J. Kaiser, Bangor University, Bangor, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction This article deals with the causes and consequences of habitat modification in the marine environment, that occur as a result of human activities. The effects of natural forcing (waves, currents, glacial scour, seismic activity, and natural climate change) and how they interact with habitats are dealt with elsewhere. Nevertheless, in order to place the effects of human activities on habitats in a meaningful context, it is necessary to appreciate the different scales at which natural and human-mediated processes occur. Habitat modifications through natural processes can occur as a result of direct physical or chemical processes (e.g., alteration of temperature, seismic activity, and input of freshwater) or through biologically mediated processes (e.g., changes in species composition or dominance). At the largest scale, marine habitats are influenced by the interaction between atmospheric changes and the water column, which in turn affect the biological processes that occur in the habitat. These changes in the watercolumn environment occur over the largest spatial scales (entire oceans) and have periodicities that vary in timescale from millennia to decades (e.g., the North Atlantic Oscillation (NAO)).
Large-Scale Processes: Natural Forcing The NAO is a good example of how physical processes can affect habitat. Decadal changes in the NAO affect wind forcing and precipitation levels according to whether the NAO index is positive or negative. In years when the index is positive, average wind speeds are higher leading to greater physical mixing of the water column, which will affect the timing and persistence of stratification on the continental shelf Northeast Atlantic. When in a negative phase, lower wind speeds are associated with an increased incidence of anoxic events due to prolonged periods of stratification in coastal waters. The latter is particularly pronounced in enclosed bodies of water such as the Baltic Sea. Wind forcing also generates stress at the seabed in shallow water near the coast. These
areas are among the most biologically productive and support important nursery habitats for commercially important fish, and diving and wading birds. Wave stress is a direct controlling factor on benthic production which increases with depth and distance from the shore to a point where it reaches a maximum before decreasing again. Increasing levels of wind stress will decrease near-shore benthic production and lower coastal-carrying capacity for species dependent upon the benthos (Figure 1). Whenever the NAO is positive, increasing precipitation will elevate the amount of sediment discharged from rivers and is associated with increased frequencies of extreme sediment-discharge events linked with flash floods. Such events in large river systems (e.g., the River Amazon) can cause mass mortalities of benthos 10 km2 offshore. Elevated inputs of freshwater discharge will affect the extent of density-driven coastal currents in regions of freshwater influence (ROFIs) leading to extended front systems that run in parallel with the coastline. In addition to these periodic changes in the oceanic regime, we can expect global climate change to greatly exacerbate the extreme nature of many aspects of physical forcing that influences habitat modification in marine systems.
Smaller-Scale Processes: Natural and Biological Forcing The examples given above operate at large scales and influence both water-column and seabed habitat properties. Other more localized natural forcing events can have profound impacts upon the habitat. Examples would include glacial scour that can remove entire habitats on a scale of a few kilometers, localized seismic activity resulting in gas discharge or geothermal activity, localized coastal erosion and daily tidal scour, and carrion falls to the deep seafloor. Apart from the last example these individual processes operate at much larger scales than the modifying processes undertaken by individual biota. Such processes include grazing, predation, and bioturbation. Despite the relatively small scale of individual habitat-modifying events (e.g., an intertidal fiddler crab excavating a burrow), the sum total of many individuals events can equate to a significant disturbance process. However, natural forcing can lead to changes in biological community structure, and habitat changes can arise if ecosystem engineering organisms are affected. Ecosystem engineers
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Figure 1 (a) The distribution of shellfish eating diving sea ducks in Liverpool Bay, UK, with mean water depth, (b) the relationship between wave stress and depth, and (c) the relationship between bivalve biomass with depth and wave stress in the same area. Shell flat is more exposed to prevailing winds than the sheltered North Wales coast. Consequently, a peak in bivalve biomass occurs in shallower water off North Wales but in deeper water (where wave stress is lower) on shell flat. The distribution of the diving sea ducks corresponds reasonably well with the observed peak in biomass when it is considered that diving depth limits prey availability. Adapted from Kaiser MJ, Clarke KR, Hinz H, Austen MCV, Somerfield PJ, and Karakassis I (2006) Global analysis and prediction of the response of benthic biota and habitats to fishing. Marine Ecology Progress Series 311: 1–14.
are those taxa that have a strong influence on some component of a marine habitat. A good example of the latter is the fluctuation between kelp and sea urchin-dominated systems in response to periodic natural outbreaks of urchin disease (natural forcing). When urchin population biomass is high, kelp biomass is grazed down, but the latter are released from grazing when urchin populations are reduced by disease. Such systems are also affected by human activities and we will return to this later.
Habitat Modification as an Ecological Process In summary, natural forcing provides a background of habitat modification at generally large scales against which are caste the smaller-scale, more localized natural forcing events and biologically mediated modifying activities. Habitat modification creates disturbance which is a critical process in the maintenance of alpha and beta diversity in terrestrial and aquatic systems. Consequently, the seafloor is a patchwork of communities in different stages of recovery, succession, or climax. It is against this natural background of disturbance and habitat
modification that the impact and significance of human activities needs to be assessed.
Human Forcing The effects of human activities can be split broadly into: direct physical effects on the habitat, and indirect effects through alteration of habitat quality or through changes in biological composition. Direct effects are usually more immediate in terms of their impact and are more readily quantifiable (e.g., aggregate extraction, fishing disturbance, and land reclamation), whereas indirect effects may remain unrecognized for many years or until they interact with some other stressor to induce a habitat change (e.g., increasing levels of sublethal pollutants). Human activities are likely to induce long-lasting habitat changes, if the scale, frequency, or intensity of the human activity exceeds the background natural disturbance regime.
Fishing One of the most widespread agents of direct habitat modification is the use of fishing gears that are towed
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HABITAT MODIFICATION
Y = loge (1 + (% change from control)/101)
across the seabed. This source of habitat disturbance has been the subject of over 100 experimental manipulations conducted over the last 25 years. In general, as fishing gear is towed across the seabed, parts of the gear dig into the substratum. This causes resuspension of sediments into the water column and, in deeper mud areas, can lead to significant transport of sediment down the continental shelf slope. The consequences of sediment resuspension in more dynamic habitats are likely to be less significant; however, this remains a much understudied topic. What is unequivocal is the direct mortality caused by the direct physical contact of fishing gear with benthic biota. The use of meta-analytic approaches has enabled the general response of benthic biota to different fishing gears in different habitats. What emerges from this analysis is that the direct mortality caused varies among different fishing gears and different habitats. While this result is not surprising, it has important management implications. Scallop (Pectinid) dredges have the most negative effect of all fishing gears across all habitat types, whereas the response of biota to beam trawls varies considerably with habitat type (Figure 2). Perhaps of greater importance is the time taken for disturbed communities to recover to a condition comparable with similar unimpacted areas. These empirical studies have formed the basis for a body-size-based model of the response of benthic communities to fishing disturbance which was validated extensively
with field data. This model has been used to demonstrate that the response of benthic biomass and production to fishing disturbance follows a negative power function such that relatively low levels of fishing intensity cause the greatest proportional loss of virgin biomass and production. Consequently, benthic communities in areas that are fished 2 or 3 times per annum differ little in absolute terms from similar benthic communities that are fished perhaps 10 or more times per annum. This finding is also applicable to other forms of chronic physical disturbance. The preceding observations and predictions apply to seabed habitats on the continental shelf. As yet no one has undertaken experimental manipulations of fishing disturbance in the deep sea. Nevertheless, the life history of deep-sea biota is such that the consequences of physical impacts by fishing gear on the deep-sea benthos and associated habitats will be measured in terms of years to hundreds of years. Recent discoveries of cold-water coral formations on the continental slope in the North Atlantic have coincided with clear evidence of the direct destruction of some of these structures by towed demersal trawl gear. The sensitivity and vulnerability of such longlived biogenic structures subsequently led to the creation of fishing-exclusion areas designed to protect these habitats. Fishing can also modify habitat indirectly through changes in trophic interactions. This is most likely to
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Figure 2 The mean initial response (up to 7 days after impact) of deposit- and suspension-feeding fauna to beam trawling (BT), otter trawling (OT), and scallop dredging (SD) in gravel (G), sand (S), and muddy sand/mud habitats combined (M). Also shown are 95% confidence intervals (dashed lines indicate only two points available for the mean calculation and hence intervals usually extending outside the plotted range). Values above the x-axis denote the number of data points in each mean calculation. An adequate test for a significant initial impact is to note whether the 95% confidence interval crosses the zero-response line (central dashed line). The key point to note from this figure is how the initial response of the biota to different fishing gears varies to a different extent among habitat types. Thus, the response to beam trawling is very variable, whereas scallop dredging appears to have a consistently negative effect in all habitats. Adapted from Kaiser MJ, Galanidi M, Showler DA, et al. (2006) Distribution and behaviour of common scoter Melanitta nigra relative to prey resources and environmental parameters. Ibis 148: 110–128.
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occur in systems in which keystone species exist that exert a strong influence on other species at a higher or lower trophic level to themselves. Such systems are prone to ‘trophic cascades’. A good example of the latter is the relationship between fishing, sea lions, killer whales, sea otters, sea urchins, and kelp on the west coast of Canada. Fishing activities have depleted populations of forage fish that supported the sea lion population. As sea lions are an important prey species for killer whales, the latter have begun to predate sea otters which in turn feed on sea urchins. The reduction in sea otter population size has released sea urchins from predation pressure such that they have increased in abundance and have grazed down kelp beds such that only urchin barrens dominated by small fleshy algae remain. Similar cascade effects have also been reported for coral reef systems. Cascade effects are manifested relatively quickly due to the strong linkage between each of the components in the system. In other more complex food webs in which the linkages among different species are more complex, the entire removal of one species may have little or no effect on other components. The reason for this lack of response is that multiple species at each trophic level may utilize the same resources; thus, any resources made available after the demise of one species will be utilized by another. This is a phenomenon termed ‘diffuse predation’. Thus, ecosystem complexity provides a buffer against the effects of harvesting specific species. However, a current pressing challenge is to be able to understand when we risk overstepping the threshold beyond which sufficient components of the ecosystem have been removed to alter its functioning. The latter has been thought to occur off the eastern coast of Canada. The removal of demersal fish biomass coupled with a reduction in benthic biomass has resulted in a breakdown in benthopelagic coupling. Consequently, the system has undergone a phase shift such that most of the energy within this ecosystem is recycled within the pelagic compartment.
Aquaculture Wild-capture fishery landings peaked in the early 1990s, while the global contribution of aquaculture to the harvest of marine species has continued to increase at a rate of 10% per year since 1990. Most mariculture practices occur within 10 nautical miles of the coastline due to logistical and technological challenges. Habitat modification can occur through a change in land use through the deforestation of
mangroves for the construction of shrimp ponds or through changes in hydrography and inputs of organic matter by suspended cultivation systems. Most of the important habitat modification processes related to aquaculture relate to pollution from excessive organic inputs from excessive feed and the production of feces. The resulting elevated levels of organic input lead to classic responses in the benthic fauna (Figure 2). Extreme cases of organic-enrichment deoxygenation of the water column and sediment can result in anoxic sediments in the immediate area of cultivation and a benthic community dominated by mats of bacteria. In addition to organic enrichment, the increasing use of antimicrobial products and chemical agents to control sea-lice can alter the microbial composition of the sediment in the immediate vicinity of suspended cages used for fish cultivation. In general, however, the effects of cage cultivation on sediment habitats are relatively localized and can be alleviated by rotating the use of different areas such that areas of the seabed are allowed to recover for a period of time. In enclosed bodies of water, it is possible to modify the water-column habitat by overstocking the biomass of cultivated species such that the carrying capacity of the system is exceeded. Examples of this have occurred in enclosed bays in France (Baie D’Archacon), where oyster culture was allowed to develop in an uncontrolled manner. Overstocking led to poor growth rates (due to competition for food) and changes in benthic community composition. Some bivalve culture systems use direct laying of the stock onto the seabed. This can induce hydrological changes directly above the bivalve bed (e.g., in the case of cockles and mussels) such that turbulence increases. This has the effect of increasing the advection rate of phytoplankton and other particles down to the seabed (and hence enhances food supply) but can also increase the sedimentation rate. These systems are often typified by raised areas of the seabed that are occasionally leveled by the cultivator. The associated community changes that occur with on-bed cultivation are often dramatic, but are confined to a limited area within and around the cultivated plots (Figure 3). Overdevelopment of shrimp culture ponds along the coastal margin has frequently led to degradation of water quality and has exacerbated the spread of disease such that the cultivation activities have become unviable. These environmental problems are well understood and new developments typically involve the use of cultivation practices that are better integrated into natural systems to alleviate and prevent their potential negative effects (e.g., by planting additional mangroves).
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HABITAT MODIFICATION
Distant control plots (no mussels) Number of species
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ln(n + 1) Mussel area cm−2 Figure 3 The response of species richness to the mussel cover in commercially cultivated plots of mussels. At very low mussel cover there is a suggestion that species richness may increase slightly as indicated by the dashed line; however, increasing the extent of mussel cover is associated with a sharp decline in species richness. Nevertheless, these effects are highly restricted to the area of cultivation as indicated by reference to ‘distant control plots’ with no mussel cover. Adapted from Beadman HA, Kaiser MJ, Galanidi M, Shucksmith R, and Willows RI (2004) Changes in species richness with stocking density of marine bivalves. Journal of Applied Ecology 41: 464–475.
Aggregate Extraction and Mining The ecological effects of mining of minerals and aggregates are similar to fishing to the extent that both processes involve the direct removal of the habitat from the seabed. Compared to the spatial area affected by towed bottom fishing gear, the effects of mining and dredging are confined to smaller areas. Nevertheless, the intensity of the impact is much greater (i.e., the depth to which the habitat is affected). Aggregates are often stable substrata that are not easily reworked by natural physical forcing, while mining can occur on the abyssal plain where physical forcing processes are much reduced. Consequently, recovery rates for both the habitats and their associated fauna are likely to exceed those that occur in response to bottom fishing in the same habitat.
River Discharge: Quality and Quantity Agricultural practices such as deforestation have led to a fourfold increase in nutrient loads in river discharge and are associated with phytoplankton blooms that generate mainly diatomaceous phytodetritus. This organic material fuels rapid rates of microbial community respiration and results in depletion of oxygen at the sediment–water interface. Such events are usually periodic and tend to coincide with periods of calm weather when inshore waters undergo stratification. The precise timing and extent of such events will vary interannually according to
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the exact climatic conditions that prevail and upon the quality and quantity of river discharge. Enclosed water basins such as the Baltic Sea, the Adriatic, and the Gulf of Mexico are well-documented examples of locations where such events occur on a regular basis. These systems are characterized by restricted tidal inundation, poor water exchange, and are highly susceptible to stratification. While anoxic events would appear to be natural events associated with areas of river discharge, the extent and severity of these events would appear to be related to human activities. Nitrogen loading doubled in the River Mississippi catchment between the 1900s and 1980s resulting in eutrophication that has increased the scale of anoxic events in the Gulf of Mexico. This supposition is supported by studies of stable isotope signatures and organic tracers in sediment cores which indicate that nutrient levels started to increase in the 1950s and finally leveled off in the 1980s, and can be linked to the concomitant rise in the use of artificial fertilizers. In addition to concern regarding the effects of elevated nutrient inputs into coastal waters, there are many other contaminants that persist in the marine environment, particularly in sediments. Many of these contaminants, such as radionuclides and organic substances, are derived from industrial and heat-generation sources. Polychlorinated biphenols (PCBs) are particularly persistent and accumulate through the food chain with negative effects expressed in predators at the top of food chains. While the effects of such contaminants on apex predators have been documented in a number of cases, the effects on organisms at lower trophic levels are less well studied. Although invertebrates such as amphipods are used in lethal toxicity tests, the incidence of contaminant mortality associated with bioengineering organisms that affect habitat structure remains unknown. The effects of contaminants may be subtle in that they do not necessarily cause direct mortality but may have negative effects on recruitment processes and larval viability. Sublethal effects may lower the cellular energy activity of organisms such that they are no longer able to cope with more stressful environments. Thus there is a distinct possibility that a combination of factors (increasing temperature, contaminant load) could have lethal effects. The damming of rivers for the purposes of water abstraction has severely lowered the discharge of some river systems with negative consequences for local ecosystems. For example, in the Adriatic Sea, output of the River Po and adjacent river systems has been lowered by 12% in the recent years. The reduction in nutrient inputs has been associated with a
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decrease in primary production in the local shelf water mass. The structure of many river delta regions is maintained by the supply of suspended river-born sediment. However, human interference with the supply of suspended sediment can lead to alteration in the structure of deltaic regions. For example, the construction of the barrage on the River Nile in 1868 led to the reversal of the accretion of the Nile delta, a situation later worsened by the construction of the Aswan dam. This led to coastal erosion rates of between 5 and 240 m per year. The associated reduction in nutrients transported out to sea was linked with a decline in the catches of sardines. In later years, fishery production increased as the reductions in nutrients supplied through the Nile discharge were replaced by increasing coastal urbanization.
Petroleum Exploration and Production The impacts of oil production on shelf sea ecosystems have been well reviewed. The ecological effects of the contaminants of drilling muds typically lead to a reduction in species diversity, the effects of which ameliorate with increasing distance from the drilling platform. In the North Sea, the chronic use of drilling muds over a period of 6–9 years was associated with community effects at a distance of up to 6 km from the drilling platform, although in cases when water-based muds were used, the ecological effects were reduced. In other areas, the effects may be much more localized as in the case of gas platforms in the Gulf of Mexico where the ecological effects of the drilling muds were confined to a distance of 100 m from the drilling platform. Thus, as with fishing, the impact of hydrocarbon exploration depends upon the nature of the habitat that is affected by this activity.
Discussion The extent of natural and human agents of habitat modification is highly varied both in terms of the temporal and spatial scale at which they affect habitats. Smaller-scale processes should be viewed as being nested with the larger-scale forcing processes that affect global or basin-scale processes. Direct mitigation of human activities that have negative effects on habitat structure that lead to phase shifts or long-term change (such as fishing) are amenable to management at a national or local level. Understanding the interaction between a particular human
activity and the response of different habitat types is a necessary prerequisite for effective spatial management that can lead to sustainable use and conservation of the marine environment.
See also Breaking Waves and Near-Surface Turbulence. Cold-Water Coral Reefs. Groundwater Flow to the Coastal Ocean. Hypoxia. Mangroves. Mariculture Overview.
Further Reading Beadman HA, Kaiser MJ, Galanidi M, Shucksmith R, and Willows RI (2004) Changes in species richness with stocking density of marine bivalves. Journal of Applied Ecology 41: 464--475. Estes JA and Duggins DO (1995) Sea otters and kelp forests in Alaska: Generality and variation in a community ecology paradigm. Ecological Monographs 65: 75--100. Gray JS, Clarke KR, Warwick RM, and Hobbs G (1990) Detection of initial effects of pollution on marine benthos: An examination from the Ekofisk and Eldfisk oilfields, North Sea. Marine Ecology Progress Series 66: 285--299. Hall SJ (1994) Physical disturbance and marine benthic communities: Life in unconsolidated sediments. Oceanography and Marine Biology Annual Review 32: 179--239. Jennings and Kaiser M (1998) The effects of fishing on marine ecosystems. Advances in Marine Biology 34: 201--352. Kaiser MJ, Clarke KR, Hinz H, Austen MCV, Somerfield PJ , and Karakassis I (2006) Global analysis and prediction of the response of benthic biota and habitats to fishing. Marine Ecology Progress Series 311: 1--14. Kaiser MJ, Galanidi M, Showler DA, et al. (2006) Distribution and behaviour of common scoter Melanitta nigra relative to prey resources and environmental parameters. Ibis 148: 110--128. Olsgard F and Gray JF (1995) A comprehensive analysis of the effects of offshore oil and gas exploration and production on the benthic communities of the Norwegian continental shelf. Marine Ecology Progress Series 122: 277--306. Pearson TH and Rosenberg R (1987) Macrobenthic succession in relation to organic enrichment and pollution of the marine environment. Oceanography and Marine Biology Annual Review 16: 229--311. Planques A, Guilen J, and Puig P (2001) Impact of bottom trawling on water turbidity and muddy sediment of an unfished continental shelf. Limnology and Oceanography 46: 1100--1110.
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HEAT AND MOMENTUM FLUXES AT THE SEA SURFACE P. K. Taylor, Southampton Oceanography Centre, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1187–1195, & 2001, Elsevier Ltd.
Introduction The maintenance of the earth’s climate depends on a balance between the absorption of heat from the sun and the loss of heat through radiative cooling to space. For each 100 W of the sun’s radiative energy entering the atmosphere nearly 40 W is absorbed by the ocean – about twice that adsorbed in the atmosphere and three times that falling on land surfaces. Much of this oceanic heat is transferred back to the atmosphere by the local sea to air heat flux. The geographical variation of this atmospheric heating drives the weather systems and their associated winds. The wind transfers momentum to the sea causing waves and the wind-driven currents. Major ocean currents transport heat polewards and at higher latitudes the sea to air heat flux significantly ameliorates the climate. Thus the heat and momentum fluxes through the ocean surface form a crucial component of the earth’s climate system. The total heat transfer through the ocean surface, the net heat flux, is a combination of several components. The heat from the sun is the short-wave radiative flux (wavelength 0.3–3 mm). Around noon on a sunny day this flux may reach about 1000 W m2 but, when averaged over 24 h, a typical value is 100–300 W m2 varying with latitude and season. Part of this flux is reflected from the sea surface – about 6% depending on the solar elevation and the sea state. Most of the remaining short-wave flux is absorbed in the upper few meters of the ocean. In calm weather, with winds less than about 3 m s1, a shallow layer may be formed during the day in which the sea is warmed by a few degrees Celsius (a ‘diurnal thermocline’). However, under stronger winds or at night the absorbed heat becomes mixed down through several tens of metres. Thus, in contrast to land areas, the typical day to night variation in sea surface sea and air temperatures is small, o11C. Both the sea and the sky emit and absorb long-wave radiative energy (wavelength 3–50 mm).
Because, under most circumstances, the radiative temperature of the sky is colder than that of the sea, the downward long-wave flux is usually smaller than the upward flux. Hence the net long-wave flux acts to cool the surface, typically by 30–80 W m2 depending on cloud cover. The turbulent fluxes of sensible and latent heat also typically transfer heat from sea to air. The sensible heat flux is the transfer of heat caused by difference in temperature between the sea and the air. Over much of the ocean this flux cools the sea by perhaps 10–20 W m2. However, where cold wintertime continental air flows over warm ocean currents, for example the Gulf Stream region off the eastern seaboard of North America, the sensible heat flux may reach 100 W m2. Conversely warm winds blowing over a colder ocean region may result in a small sensible heat flux into the ocean – a frequent occurrence over the summertime North Pacific Ocean. The evaporation of water vapor from the sea surface causes the latent heat flux. This is the latent heat of vaporization which is carried by the water vapor and only released to warm the atmosphere when the vapor condenses to form clouds. Usually this flux is significantly greater than the sensible heat flux, being on average 100 W m2 or more over large areas of the ocean. Over regions such as the Gulf Stream latent heat fluxes of several hundred W m2 are observed. In foggy conditions with the air warmer than the sea, the latent heat flux can transfer heat from air to sea. In summertime over the infamous fog-shrouded Grand Banks off Newfoundland the mean monthly latent heat transfer is directed into the ocean, but this is an exceptional case.
Measuring the Fluxes The standard instruments for determining the radiative fluxes measure the voltage generated by a thermopile which is exposed to the incident radiation. Typically the incoming short-wave radiation is measured by a pyranometer which is mounted in gimbals for use on a ship or buoy (Figure 1). For better accuracy the direct and scattered components should be determined separately but, apart from at the Baseline Surface Radiation Network stations which are predominantly situated on land, at present this is rarely done. The reflected short-wave
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Figure 1 A pyranometer used for measuring short-wave radiation. The thermopile is covered by two transparent domes. (Photograph courtesy of Southampton Oceanography Centre.)
radiation is normally determined from the sun’s elevation and lookup tables based on the results of previous experiments. The pyrgeometer used to determine the long-wave radiation is similar to the pyranometer but uses a coated dome to filter out, as far as possible, the effects of the short-wave heating. Because the air close to the sea surface is normally near to the sea temperature, the use of gimbals is less important. However, a clear sky view is required and a number of correction terms have to be calculated for the temperature of the dome and any short-wave leakage. Again, only the downward component is normally measured; the upwards component is calculated from knowledge of the sea temperature and emissivity of the sea surface. The turbulent fluxes may be measured in the nearsurface atmosphere using the eddy correlation method. If upward moving air in an eddy is on average warmer and moister than the downward moving air, then there is an upwards flux of sensible heat and water vapor and hence also an upward latent heat flux. Similarly the momentum flux, or wind stress, may be determined from the correlation between the horizontal and vertical wind fluctuations. Since a large range of eddy sizes may contribute to the flux, fast response sensors capable of sampling at 10 Hz or more must be exposed for periods of the order of 30 min for each flux determination. Three-component ultrasonic anemometers (Figure 2) are relatively robust and, by also determining the speed of sound, can provide an estimate of the sonic temperature flux, a function of the heat and moisture fluxes. The sensors used for determining the fluctuations in temperature and humidity have previously tended to be fragile and prone to contamination by salt particles which are ever-present in the marine atmosphere. However,
Figure 2 The sensing head of a three-component ultrasonic anemometer. The wind components are determined from the different times taken for sound pulses to travel in either direction between the six ceramic transducers. (Photograph courtesy of Southampton Oceanography Centre.)
improved sonic thermometry, and new techniques for water vapor measurement, such as microwave refractometry or differential infrared absorption instruments, are now becoming available. Despite these improvements in instrumentation, obtaining accurate eddy correlation measurements over the sea remains very difficult. If the instrumentation is mounted on a buoy or ship the six components of the wave-induced motion of the measurement platform must be measured and removed from the signal. The distortion both of the turbulence and the mean wind by ship, buoy or fixed tower must be minimized and, as far as possible, corrected for. Thus eddy correlation measurements are not routinely obtained over the ocean, rather they are used in special air–sea interaction experiments to calibrate other less direct methods of flux estimation. For example, in the inertial dissipation method, fluctuations of the wind, temperature, or humidity at a few Hertz are measured and related (through turbulence theory) to the fluxes. This method is less sensitive to flow distortion or platform motion, but relies on various assumptions about the formation and dissipation of turbulent quantities, which may not be valid under some conditions. It has been implemented
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HEAT AND MOMENTUM FLUXES AT THE SEA SURFACE
on a semi-routine basis on some research ships to increase the range of available flux data. The most commonly used method of flux estimation is variously referred to as the bulk (aerodynamic) formulae. These formulae relate the difference between the value of temperature, humidity or wind (‘x’ in [1]) at some measurement height, z, and the value assumed to exist at the sea surface – respectively the sea surface temperature, 98% saturation humidity (to allow for salinity effects), and zero wind (or any nonwind-induced water current). Thus the flux Fx of some quantity x is: Fx ¼ rUz Cxz ðxz x0 Þ
½1
where r is the air density, and Uz the wind speed at the measurement height. While appearing intuitively correct (for example, blowing over a hot drink will cool it faster) these formulae can also be derived from turbulence theory. The value for the transfer coefficient, Cxz, characterizes both the surface roughness applicable to x and the relationship between Fx and the vertical profile of x. This varies with the atmospheric stability, which itself depends on the momentum, sensible heat, and water vapor fluxes, as well as the measurement height. Thus, although it may appear simple, Eqn [1] must be solved by iteration, initialized using the equivalent neutral value of Cxz at some standard height (normally 10 m), Cx10n. Typical neutral values (determined using eddy correlation or inertial dissipation data) are shown in Table 1. Many research problems remain. For example: CD10n is expected to depend on the state of development of the wave field, but can this be successfully characterized by the ratio of the predominant wave speed to the wind speed (the wave age), or by the wave height and steepness, or is a spectral representation of the wave field required? What are the effects of waves propagating from other regions (i.e., swell waves)? What is the behavior of CD10n in low wind speed conditions? Furthermore CE10n and CH10n are relatively poorly defined by the available experimental data, and Table 1
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recent bulk algorithms have used theoretical models of the ocean surface (known as surface renewal theory) to predict these quantities from the momentum roughness length.
Sources of Flux Data Until recent years the only source of data for flux calculation routinely available from widespread regions of the world’s oceans was the weather reports from merchant ships. Organized as part of the World Weather Watch system of the World Meteorological Organisation, these ‘Voluntary Observing Ships (VOS)’ are asked to return coded weather messages at 00 00, 06 00, 12 00, and 18 00 h GMT daily, also recording the observation (with further details) in the ship’s weather logbook. The very basic set of instruments provided will normally include a barometer and a means of measuring air temperature and humidity – typically wet and dry bulb thermometers mounted in a hand swung sling psychrometer or a fixed, louvered ‘Stevenson’ screen. Sea temperature is obtained using a thermometer and an insulated bucket, or by reading the temperature gauge for the engine cooling water intake. Depending on which country recruited the VOS an anemometer and wind vane might be provided, or the ship’s officers might be asked to estimate the wind velocity from observations of the sea state using a tabulated ‘Beaufort scale’. Because of the problems of adequately siting an anemometer and maintaining its calibration, these visual estimates are not necessarily inferior to anemometer-based values. Thus the VOS weather reports include all the variables needed for calculating the turbulent fluxes using the bulk formulae. However, in many cases the accuracy of the data is limited both by the instrumentation and its siting. In particular, a large ship can induce significant changes in the local temperature and wind flow, since the VOS are not equipped with radiometers. The short-wave and long-wave fluxes must be estimated from the observer’s estimate
Typical values (with estimated uncertainties) for the transfer coefficientsa
Flux
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Typical values
Momentum
Drag coefficient CD10n( 1000) Stanton no., UH10n Dalton no., UE10n
¼ 0.61 (70.05) þ 0.063 (70.005) U10n (U10n43 m s1) ¼ 0.61 þ 0.57/U10no3 m s1 1.1 (70.2) 103 1.2 (70.1) 103
Sensible heat Latent heat
a Neither the low wind speed formula for CD10n, nor the wind speed below which it should be applied, are well defined by the available, very scattered, experimental data. It should be taken simply as an indication that, at low wind speeds, the surface roughness increases as the wind speed decreases due to the dominance of viscous effects.
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of the cloud amount plus (as appropriate) the solar elevation, or the sea and air temperature and humidity. The unavoidable observational errors and the crude form of the radiative flux formulae imply that large numbers of reports are needed, and correction schemes must be applied, before satisfactory flux estimates can be obtained. While there are presently nearly 7000 VOS, the ships tend to be concentrated in the main shipping lanes. Thus whilst coverage in most of the North Atlantic and North Pacific is adequate to provide monthly mean flux values, elsewhere data is mainly restricted to relatively narrow, major trade routes. For most of the southern hemisphere the VOS data is only capable of providing useful values if averaged over several years, and reports from the Southern Ocean are very few indeed. These shortcomings of VOS-derived fluxes must be borne in mind when studying the flux distribution maps presented below. Satellite-borne sensors offer the potential to overcome these sampling problems. They are of two types, passive sensors which measure the radiation emitted from the sea surface and the intervening atmosphere at visible, infrared, or microwave frequencies, and active sensors which transmit microwave radiation and measure the returned signal. Unfortunately these remotely sensed data do not allow all of the flux components to be adequately estimated. Sea surface temperature has been routinely determined using visible and infrared radiometers since about 1980. Potential errors due, for example, to changes in atmospheric aerosols following volcanic eruptions, mean that these data must be continually checked against ship and buoy data. Algorithms have been developed to estimate the net surface short-wave radiation from top of the atmosphere values; those for estimating the net surface long wave are less successful. The surface wind velocity can be determined to good accuracy by active scatterometer sensors by measuring the microwave radiation backscattered from the sea surface. Unfortunately scatterometers are relatively costly to operate, since they demand significant power from the spacecraft and, to date, few have been flown. The determination of near-surface air temperature and humidity from satellite is hindered by the relatively coarse vertical resolution of the retrieved data. A problem is that the radiation emitted by the near-surface air is dominated by that originating from the sea surface. Statistically based algorithms for determining the near-surface humidity have been successfully demonstrated. More recently neural network techniques have been applied to retrieving both air temperature and humidity; however, at present there is no routinely available product. Thus the satellite flux products for which useful accuracy has
been demonstrated are presently limited to momentum, short-wave radiation, and latent heat flux. Numerical weather prediction (NWP) models (as used in weather forecasting centers) estimate values of the air–sea fluxes as a necessary part of their calculations. Since these models assimilate most of the available data from the World Weather Watch system, including satellite data, radiosonde profiles, and surface observations, it might be expected that NWP models represent the best source of flux data. However, there are other problems. The vertical resolution of these models is relatively poor and many of the near-surface processes which affect the fluxes have to be represented in terms of larger-scale parameters. Improvements to these models are normally judged on the resulting quality of the weather forecasts, not on the accuracy of the surface fluxes; sometimes these may become worse. Indeed, the continual introduction of model changes results in time discontinuities in the output variables. This makes the determination of interannual variations difficult. Because of this, NWP centres such as the European Centre for Medium Range Weather Forecasting (ECMWF) and the US National Centers for Environmental Prediction (NCEP) have reanalyzed the past weather and have gone back several decades. The surface fluxes from these reanalyses are receiving much study. Those presently available appear less accurate than fluxes derived from VOS data in regions where there are many VOS reports; in sparsely sampled regions the model fluxes may be more accurate. There are particular weaknesses in the shortwave radiation and latent heat fluxes. New reanalyses are planned and efforts are being made to improve the flux estimates; eventually these reanalyses will provide the best source of flux data for many purposes.
Regional and Seasonal Variation of the Momentum Flux The main features of the wind regimes over the global oceans have long been recognized and descriptions are available in many books on marine meteorology (see Further Reading). The major features of the wind stress variability derived from ship observations from the period 1980–93 will be summarized here, using plots for January and July to illustrate the seasonal variation. The distribution of the heat fluxes will be discussed in the next section. In northern hemisphere winter (Figure 3A) large wind stresses due to the strong midlatitude westerly winds are obvious in the North Atlantic and the North Pacific west of Japan. To the south of these
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Figure 3 Monthly vector mean wind stress (N m2) for (A) January and (B) July calculated from Voluntary Observing Ship weather reports for the period 1980–93. (Adapted with permission from Josey SA, Kent EC and Taylor PK (1998) The Southampton Oceanography Centre (SOC) Ocean–Atmosphere Heat, Momentum and Freshwater Flux Atlas. SOC Report no. 6.)
regions the extratropical high pressure zones result in low wind stress values, south of these is the northeast trade wind belt. The Inter-Tropical Convergence Zone (ITCZ) with very light winds is close to the equator in both oceans. In the summertime southern hemisphere the south-east trade wind belt is less well marked. The extratropical high pressure regions are extensive but, despite it being summer, high winds and significant wind stress exist in the midlatitude
southern ocean. The north-east monsoon dominates the wind patterns in the Indian Ocean and the South China Sea (where it is particularly strong). The ITCZ is a diffuse region south of the equator with relatively strong south-east trade winds in the eastern Indian Ocean. In northern hemisphere summer (Figure 3B) the wind stresses in the midlatitude westerlies are very much decreased. Both the north-east and the
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south-east trade wind zones are evident respectively to the north and south of the ITCZ. This is predominantly north of the equator. The south-east trades are particularly strong in the Indian Ocean and feed into a very strong south-westerly monsoon flow in the Arabian Sea. The ship data indicate very strong winds in the Southern Ocean south west of Australia. These are also evident in satellite scatterometer data, which suggest that the winds in the Pacific sector of the Southern Ocean, while still strong, are somewhat less than those in the Indian Ocean sector. In contrast the ship data appear to show very light winds. The reason is that in wintertime there are practically no VOS observations in the far south Pacific. The analysis technique used to fill in the data gaps has, for want of other information, spread the light winds of the extratropical high pressure region farther south than is realistic; a good example of the care needed in interpreting the flux maps available in many atlases.
Regional and Seasonal Variation of the Heat Fluxes The global distribution of the mean annual net heat flux is shown in Figure 4A. The accuracy and method of determination of such flux distributions will be discussed further below, here they will be used to give a qualitative description. Averaged over the year the ocean is heated in equatorial regions and loses heat in higher latitudes, particularly in the North Atlantic. However, this mean distribution is somewhat misleading, as the plots for January (Figure 4B) and July (Figure 4C) illustrate. The ocean loses heat over most of the extratropical winter hemisphere and gains heat in the extratropical summer hemisphere and in the tropics throughout the year. The relative magnitude of the individual flux components is illustrated in Figure 5 for three representative sites in the North Atlantic Ocean. At the Gulf Stream site (Figure 5A) the large cooling in winter dominates the incoming solar radiation in the annual mean. However, even at this site the mean monthly short-wave flux in summer is greater than the cooling. Indeed the effect of the longer daylight periods increases the mean short-wave radiation to values similar to or larger than those observed in equatorial regions (Figure 5C). The midlatitude site (Figure 5B) is typical of large areas of the ocean. The ocean cools in winter and warms in summer, in each case by around 100 W m2. The annual mean flux is small – around 10 W m2 – but cannot be neglected because of the very large ocean areas involved. At this site, and generally over the ocean, this annual balance is between the sum of the latent heat flux and net
long-wave flux which cool the ocean, and the net short-wave heating. Only in very cold air flows, as over the Gulf Stream in winter, is the sensible heat flux significant. As regards the interannual variation of the surface fluxes, the major large-scale feature over the global ocean is the El Nin˜o-Southern Oscillation system in the equatorial Pacific Ocean. The changes in the net heat flux under El Nin˜o conditions are around 40 W m2 in the eastern equatorial Pacific. For extratropical and midlatitude regions the interannual variability of the summertime net heat flux is typically about 20–30 W m2, being dominated by the variations in latent heat flux. In winter the typical variability increases to about 30–40 W m2, although in particular areas (such as over the Gulf Stream) variations of up to 100 W m2 can occur. The major spatial pattern of interannual variability in the North Atlantic is known as the North Atlantic Oscillation (NAO). This represents a measure of the degree to which mobile depressions, or alternatively near stationary high pressure systems, occur in the midlatitude westerly zone.
Accuracy of Flux Estimates It has been shown that, although the individual flux components are of the order of hundreds of W m2, the net heat flux and its interannual variability over much of the world ocean is around tens of W m2. Furthermore it can be shown that a flux of 10 W m2 over 1 year would, if stored in the top 500 m of the ocean, heat that entire layer by about 0.151C. Temperature changes on a decadal time scale are at most a few tenths of a degree, so the global mean budget must balance to better than a few W m2. For these various reasons there is a need to measure the flux components, which vary on many time and space scales, to an accuracy of a few W m2. Given the available data sources and methods of determining the fluxes described in the previous sections, it is not surprising that this level of accuracy cannot be achieved at present. To take an example, in calculating the flux maps shown in Figure 4 from VOS data many corrections were applied to the VOS observations to attempt to remove biases caused by the methods of observation. For example, air temperature measurements were corrected for the heat island caused by the ship heating up in sunny, low wind conditions. The wind speeds were adjusted depending on the anemometer heights on different ships. Corrections were applied to sea temperatures calculated from engine room intake data. Despite these and other corrections, the
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HEAT AND MOMENTUM FLUXES AT THE SEA SURFACE
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Figure 4 Variation of the net heat flux over the ocean, positive values indicate heat entering the ocean: (A) annual mean, (B) January monthly mean, (C) July monthly mean. (Adapted with permission from Josey SA, Kent EC and Taylor PK (1998) The Southampton Oceanography Centre (SOC) Ocean–Atmosphere Heat, Momentum and Freshwater Flux Atlas. SOC Report no. 6.)
global annual mean flux showed about 30 W m2 excess heating of the ocean. Previous climatologies calculated from ship data had shown similar biases and the fluxes had been adjusted to remove the bias,
or to make the fluxes compatible with estimates of the meridional heat transport in the ocean. However, comparison of the unadjusted flux data with accurate data from air–sea interaction buoys showed good
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600
600
500
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0 Annual (A)
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40°N, 20°W SW radn. LW radn.
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Latent Heat Sensible Heat
400 300 200 100 0 Annual (C)
January 0°N, 20°W
July
Figure 5 Mean heat fluxes at three typical sites in the North Atlantic for the annual mean, and the January and July monthly means. In each case the left-hand column shows the fluxes which act to cool the ocean while the right-hand column shows the solar heating. (A) Gulf Stream site (401N, 601W), (B) midlatitude site (401N, 201W), (C) equatorial site (01N, 201W).
agreement between the two. This suggests that adjusting the fluxes globally is not correct and that regional flux adjustments are required; however, the exact form of these corrections is presently not shown. In the future, computer models are expected to provide a major advance in flux estimation. Recently coupled numerical models of the ocean and of the atmosphere have been run for many simulated years during which the modeled climate has not drifted. This suggests that the air–sea fluxes calculated by the models are in balance with the simulated oceanic and atmospheric heat transports. However, it does not imply that the presently estimated flux values are realistic. Errors in the short-wave and latent heat fluxes may compensate one another; indeed in a typical simulation the sea surface temperature stabilized to a value which was, over large regions of the ocean, a few degrees different from that which is observed. Nevertheless the estimation of flux values using climate or NWP models is a rapidly developing field and improvements will doubtless occur in the next few years. There will be a continued need for in situ and satellite data for assimilation into the models and for model development and verification. However, it seems very likely that in future the most accurate routine source of the air–sea flux data will
be from numerical models of the coupled ocean–atmosphere system.
See also El Nin˜o Southern Oscillation (ENSO). El Nin˜o Southern Oscillation (ENSO) Models. Evaporation and Humidity. Heat Transport and Climate. IR Radiometers. North Atlantic Oscillation (NAO). Satellite Passive-Microwave Measurements of Sea Ice. Satellite Remote Sensing of Sea Surface Temperatures. Sensors for Mean Meteorology. Turbulence Sensors. Upper Ocean Heat and Freshwater Budgets. Wave Energy. Wave Generation by Wind. Wind- and Buoyancy-Forced Upper Ocean.Wind Driven Circulation
Further Reading Browning KA and Gurney RJ (eds.) (1999) Global Energy and Water Cycles. Cambridge: Cambridge University Press. Dobson F, Hasse L, and Davis R (eds.) (1980) Air–Sea Interaction, Instruments and Methods. New York: Plenum Press.
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Kraus EB and Businger JA (1994) Atmosphere–Ocean Interaction, 2nd edn. New York: Oxford University Press. Meteorological Office (1978) Meteorology for Mariners, 3rd edn. London: HMSO.
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Stull RB (1988) An Introduction to Boundary Layer Meteorology. Dordrecht: Kluwer Academic. Wells N (1997) The Atmosphere and Ocean: A Physical Introduction, 2nd edn. London: Taylor and Francis.
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HEAT TRANSPORT AND CLIMATE H. L. Bryden, University of Southampton, Southampton, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction: The Global Heat Budget The Earth receives energy from the sun (Figure 1) principally in the form of short-wave energy (sunlight). The amount of solar radiation is quantified by the ‘solar constant’ which satellite radiometers have measured since about 1985 to have a mean value of 1366 W m 2, an 11-year sunspot cycle of amplitude 1.5 W m 2, and a maximum at maximum sunspot activity. A fraction of the sunlight is reflected directly back into space and this fraction is termed the albedo. Brighter areas like snow in polar regions have high albedo (0.8) reflecting most of the short-wave radiation back to space, while darker areas like the ocean have small albedo (0.05) and small reflection. Averaged over the Earth’s surface, the albedo is about 0.3. Overall, the net incoming short-wave radiation (incoming minus reflected) peaks in equatorial regions and decreases to small values in polar latitudes (Figure 2). The Earth radiates energy back to space in the form of long-wave, black-body radiation proportional to the fourth power of the absolute temperature at the top of the atmosphere. Because the temperature at the top of the atmosphere is relatively uniform with latitude varying only from 200 to 230 K, there is only a small latitudinal variation in outgoing radiation (Figure 1). Over a year, the net
incoming radiation equals the net outgoing radiation within our ability to measure the radiation, thus maintaining the overall heat balance of the Earth. For the radiation budget as a function of latitude, however, there is more incoming short-wave radiation at equatorial and tropical latitudes and more outgoing long-wave radiation at polar latitudes. To maintain this heating–cooling distribution, the atmosphere and ocean must transport energy poleward from the Tropics toward the Pole and the maximum poleward energy transport in each hemisphere occurs at a latitude of c. 351, where there is a change from net incoming to net outgoing radiation, and the maximum has a magnitude of about 5.8 PW (1 petawatt (PW) ¼ 1015 W). As recently as the mid-1990s, it was controversial whether the ocean or the atmosphere was responsible for the majority of the energy transport. Oceanographers found a maximum ocean heat transport of about 2 PW at 25–301 N, while meteorologists reported a maximum atmospheric transport of about 2.5 PW from the analysis of global radiosonde network observations. Thus, there was a missing petawatt of energy transport in the Northern Hemisphere for the combined ocean–atmosphere system. Recent analyses combining observations and models suggest that the atmosphere does carry the additional petawatt that observational analyses alone could not find Radiation balance Incoming radiation (1– )Q 300
300 E
E
Radiation (W m−2)
Outgoing radiation
Q
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Net radiation R
0
0 90° N 60°
Q
30°
0°
30°
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−100
−100 R = (1−)Q −E Figure 1 Schematic of the Earth’s radiation budget. Q is incoming, short-wave solar radiation; aQ is reflected solar radiation where a is albedo, and E is outgoing, long-wave black-body radiation.
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Figure 2 Latitudinal profiles of net incoming short-wave radiation, outgoing long-wave radiation, and the net radiative heating of the Earth. Note the latitudinal scale is stretched so that it is proportional to the surface area of the Earth.
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HEAT TRANSPORT AND CLIMATE
due to a sparsity of radiosonde observations over the ocean. Now it is generally accepted that the atmosphere carries the majority of energy transport at 351 N, though some oceanographers point out that half of the maximum atmospheric transport is due to latent heat (water vapor) transport that can be considered to be a joint ocean–atmosphere process (Figure 3). Here we will concentrate on ocean heat transport, especially on the mechanisms of ocean heat transport in the Northern Hemisphere. Conservation of energy in the ocean is effectively expressed as conservation of heat, where heat is defined to be rCpY, where r is seawater density, Cp is specific heat of seawater at constant pressure, and Y is potential temperature, the temperature of a water parcel brought adiabatically (without heat exchange) to the sea surface from depth. Because rCp is nearly constant at about 4.08 106 J m 3 1C 1, heat conservation is essentially expressed as conservation of potential temperature. There are many subtleties to the definitions that are described in entries for density, potential temperature, and heat, but here we use traditional definitions of potential temperature, density, and specific heat based on the internationally recognized equation of state for seawater. Ocean heat transport is then the flow of heat through the ocean, rCpYv, where v is the water
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velocity. Such definition depends on the temperature scale and has little meaning until it is considered for a given volume of the ocean. Because of mass conservation, there is no net mass transport into or out of a fixed ocean volume over long timescales (neglecting the relatively miniscule contributions from evaporation minus precipitation), so it is the ocean heat transport convergence that is meaningful, that is the amount of heat transport into the volume minus the heat transport out of the volume. Since mass is conserved, the heat transport convergence is proportional to the mass transport times the difference in temperature between the inflow and the outflow across the boundaries of the volume. For a complete latitude band like 251 N, where the Atlantic Ocean and Pacific Ocean volume north of 251 N can be considered to be an enclosed ocean, the heat transport convergence is commonly referred to as the ocean heat transport at 251 N. Individually the Atlantic and Pacific Oceans are nearly enclosed with only a small throughflow connecting them in Bering Straits, so the Atlantic heat transport at 251 N is commonly referred to, even though it is formally the heat transport convergence between 251 N and Bering Straits and similarly for Pacific heat transport at 251 N. Such definitions of heat transport are generally used throughout the Atlantic north of 301 S and the Pacific north of about 101 N, where each ocean
Components of global energy balance 6
Northward energy transport (PW)
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Dry static atmosphere Latent Ocean Total
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Figure 3 Components of atmosphere and ocean energy transports required to balance the net radiational heating/cooling of the Earth following Figure 2. The standard atmospheric energy transport is here divided into the dry static atmospheric energy transport and the latent heat transport. Latent heat transport is fundamentally a joint atmosphere–ocean process since the atmospheric water vapor transport is balanced by an opposing oceanic freshwater transport. The ocean heat transport is determined by integrating over the oceans the spatial distribution of atmosphere–surface heat exchange calculated by subtracting the atmospheric energy transport divergence from the radiative heating at the top of the atmosphere.
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basin is closed except for the small Bering Straits transport.
Air–Sea Heat Exchange Conservation of heat means that any convergence of ocean heat transport is balanced by heat loss to the atmosphere (the amount of heat gain or loss through the ocean bottom is small in comparison to exchanges with the atmosphere). Thus charts of air–sea energy exchange are primary sources for our understanding of ocean heat transport, where it occurs and how big it is. Estimates of air–sea energy exchange have long been made based on measurements of cloud cover, surface air and water temperatures, wind speed, humidity, and bulk formula exchange coefficients to calculate the size of the radiational heating and latent heat cooling of the ocean surface and the sensible heat exchange between the ocean and atmosphere. Combining such shipboard observations on a global scale produces air–sea flux climatologies giving air–sea exchange by month and region. From such climatologies, the global distribution of annual averaged air–sea energy exchange (Figure 4) shows that the ocean gains heat over much of the equatorial and tropical regions and gives up large amounts of heat over the warm poleward flowing western boundary currents like the Gulf Stream, Kuroshio or Agulhas Current, and over open-water subpolar and polar regions. Ocean heat
transport convergence for any arbitrary ocean volume can technically be estimated by summing up the air–sea energy exchange over the surface area of the ocean volume. There is a problem, however, in that the air–sea energy exchange estimates have an uncertainty of about 30 W m 2. One way to determine this uncertainty is to sum the air–sea exchanges globally and to find that there is on average a heat gain by the ocean of 30 W m 2, in each of the two state-of-the-art air–sea exchange climatologies. It is of course possible to remove this bias, either uniformly, by region or by component (radiative, latent, or sensible heat exchange); despite careful comparison with buoy measurements with bulk formula estimates at several oceanic locations, there is no consensus on how to remove the 30 W m 2 uncertainty in air–sea energy exchange.
Distribution of Ocean Heat Transport Estimating ocean heat transport convergence from in situ oceanographic measurements is a second method to quantify the role of the ocean in the global heat balance. The advantage of this direct approach is that the mechanisms of ocean heat transport are examined rather than just their overall effect in terms of the air–sea exchange. It was first reliably applied at 251 N in the Atlantic, a latitude where the warm northward Gulf Stream flow of about 30 Sv (1 Sv ¼ 1 106 m3 s 1) through Florida Straits is regularly
Net heat flux (W m−2), annual 90
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Figure 4 Global distribution of the annual mean net heat gain by the ocean as determined from bulk formula calculations. Positive values indicate a gain of heat by the oceans. SOC climatology – Josey SA, Kent EC, and Taylor PK (1999) New insights into the ocean heat budget closure problem from analysis of the SOC air–sea flux climatology. Journal of Climate 12: 2856–2880.
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HEAT TRANSPORT AND CLIMATE
monitored by submarine telephone cable voltage (the varying flow of conducting seawater through a magnetic field produces varying voltage in the cable which is continuously measured). This is the latitude of the easterly trade winds whose westward wind stress generates a northward surface water transport of about 4 Sv in the Ekman layer. Because the Atlantic is closed to the north except for a small (1 Sv) flow from the Pacific through Bering Straits, a nearly equal amount of southward flow of 35 Sv must cross the mid-ocean section between the Bahamas and Africa. The vertical distribution of this southward return flow (and its temperature) has been estimated from transatlantic hydrographic sections where temperature and salinity profiles are used to derive geostrophic velocity profiles and the reference-level velocity for the geostrophic velocity profiles is set to make the southward geostrophic transport equal to the northward Gulf Stream plus Ekman transport. Analysis of three hydrographic sections along 251 N made in 1957, 1981, and 1992 suggests that the ocean heat transport is 1.3 PW. Estimates of the error in heat transport using this direct method are about 0.3 PW, implying that direct ocean heat transport estimates are more accurate than air–sea flux estimates for areas larger than 301 latitude 301 longitude (approximately, 30 W m 2 3100 km 3100 km ¼ 0.3 PW). In terms of mechanisms, the Gulf Stream carries warm water (transport-weighted temperature of 19 1C) northward and the northward surface wind-driven Ekman transport is also warm (25 1C). The compensating southward mid-ocean flow is about equally divided between a recirculating thermocline flow above 800 m depth with an average temperature close to the 19 1C Gulf Stream flow and a cold deep-water flow with an average temperature less than 3 1C. Overall then, the northward heat transport of 1.3 PW across 251 N is due to a net northward transport of about 18 Sv of warm upper layer waters balanced by a net southward transport of cold deep waters. Deep-water formation in the Nordic and Labrador Seas of the northern Atlantic connects the northward flowing warm waters and southward flowing cold deep waters. This overall vertical circulation is commonly called the Atlantic meridional overturning circulation. During the World Ocean Circulation Experiment 1990–99, transoceanic hydrographic sections were made and ocean heat transport estimated in each ocean basin. The Atlantic heat transport is northward from 401 S to 551 N; it is of the order 0.5 PW at the southern boundary of the Atlantic, increases in the tropical regions as heat is gained at the sea surface, reaches a maximum near 251 N, and then decreases northward as the ocean gives up heat to the
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atmosphere (Figure 5). At all latitudes, the northward heat transport is due to the meridional overturning circulation where the cold deep-water transport of c. 15–20 Sv persists through the Atlantic and the compensating northward upper water flow changes temperature, warming in tropical regions and then cooling in northern subtropical and subpolar latitudes. The North Pacific, north of 101 N where the basin is essentially closed, exhibits a similar pattern of northward heat transport, but the heat transport achieves a maximum of only 0.8 PW at 251 N, 50% less than the North Atlantic despite the Pacific being more than twice as wide as the Atlantic. The Pacific is different from the Atlantic in that there is no substantial deep meridional overturning circulation, for there is no deep-water formation in the North Pacific. The Pacific heat transport is due to a horizontal circulation where warm waters flow northward in the Kuroshio western boundary current and then recirculate southward over the vast zonal extent of the Pacific, all at depths shallower than 1000 m. The Kuroshio has about the same size and temperature structure as the Gulf Stream, but the zonal temperature distribution in the Pacific exhibits much colder upper water temperatures in mid-ocean and particularly along the eastern boundary of western North America than does the Atlantic along Europe and Africa. Thus, the northward heat transport in the North Pacific is due to the horizontal upper water circulation where warm water flows northward in the Kuroshio, loses heat to the atmosphere at latitudes south of 501 N, and then recirculates southward at colder temperatures in the mid- and eastern Pacific. South of 101 N in the Pacific and throughout the Indian Ocean, ocean heat transports are somewhat ambiguous because of the throughflow from the Pacific to the Indian Ocean through the Indonesian archipelago. The throughflow transport is not yet well defined but there is a substantial observational effort now underway to quantify its mass transport and associated temperature structure. Because mass is not conserved for individual Pacific or Indian zonal sections south of 101 N, heat transports across such zonal sections in the literature generally depend on the temperature scale used as well as on the size of the throughflow assumed. These fluxes should properly be called temperature transports and to be interpreted properly should include both a net mass transport across each section and a transport-weighted temperature. Reported northward temperature transports in the South Pacific are mainly due to a net northward transport (to balance the throughflow) times temperature in degrees Celsius. Similarly, southward temperature transports in the south Indian Ocean
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60° N 0.6
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Figure 5 Heat transport and temperature transport across hydrographic sections taken during the World Ocean Circulation Experiment. For the Atlantic north of 401 S and the Pacific north of 101 N, heat transport values are presented where there is no net mass transport across the sections. For the remaining sections, temperature transports are presented where there is a net mass transport across the section and the temperature transport includes the net transport multiplied by temperature in degrees Celsius.
tend to be large because they include a substantial net southward mass transport (to balance the throughflow) times an average temperature. Careful ocean heat transport divergence estimates for the Indian Ocean taking into account the throughflow transport and temperature generally suggest that the Indian Ocean gains heat from the atmosphere between 0.4 and 1.2 PW depending on the size and temperature of the throughflow. For the South Pacific, the ocean heat transport convergence or divergence is ambiguous for a normal range of throughflow transport, so it is uncertain whether the South Pacific as a whole gains heat or loses heat to the atmosphere. It is possible to combine the South Pacific and Indian sections to enclose a confined ocean volume north of say 321 S. For such combinations where mass is conserved, the Indo-Pacific Ocean heat transport at 321 S is meaningful and estimates are that it is southward with a magnitude of 0.4–1.2 PW. For the complete latitude band at about 301 S combining Indo-Pacific and Atlantic Ocean heat transports, the ocean heat transport is southward or
poleward but with only a maximum value of about 0.5 PW since the northward Atlantic heat transport partially cancels the southward Indo-Pacific heat transport. Thus, it appears that the ocean contributes much less than the atmosphere to the poleward heat transport in the Southern Hemisphere required to balance the Earth’s radiation budget.
Eddy Heat Transport In addition to heat transport by the steady ocean circulation, temporally and spatially varying currents with scales of 10–100 days and 10–100 km, which are called mesoscale features or eddies, can also transport heat. Correlations between time-varying velocity and temperature fluctuations can transport heat, rCp /Y0 v0 S (where primes denote fluctuations and angular brackets indicate time averages), even when there is no net velocity or mass transport, that is, /v0 S is zero. Such eddy heat transport can be substantial in regions where there are strong lateral temperature gradients like the Gulf Stream or Kuroshio extensions separating
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the subtropical from the subpolar gyre. In the zonally unbounded Southern Ocean where there are several thermal fronts associated with the Antarctic Circumpolar Current, eddy heat transport is the dominant mechanism for transporting heat poleward. Here eddy motions are observed to have colder temperature when they are flowing northward and warmer temperature when following southward though there is no average flow over the eddy scales. The resulting eddy heat transport, typically of the order 4 103 W m 2 for /Y0 v0 S ¼ 0.1 1C cm s 1, is southward, downgradient from high temperature on the northern side toward cold temperature on the southern side of the front, and this downgradient heat flux is a signature of the baroclinic instability process by which eddies form and grow on the potential energy stored in the large-scale lateral temperature distribution. For the 3500 m depth and 20 000 km zonal extent of the Southern Ocean, this poleward eddy heat flux amounts to 0.3 PW across a latitude of 601 S. Individual eddies or rings of isolated water mass properties may also transport heat. For example, Agulhas rings formed with a core of Indian Ocean water properties in the retroflection area south of Africa are observed to transit across the South Atlantic. These eddies have relatively warm water cores and their heat transport is often estimated by multiplying their heat content anomaly by an estimated number of how many such rings are formed each year. Such calculation is somewhat ambiguous because it is not clear how the mass is returned and what its temperature is. Similar estimates have been made with Gulf Stream and Kuroshio rings, both warm core and cold core, and with meddies formed from the outflow of Mediterranean water. While individual rings are impressive, it is not yet clear whether they carry a significant amount of heat compared with the annual averaged air–sea exchange in any region.
Future Developments There is a third method, the residual method, for estimating ocean heat transport that takes the difference between the energy transport required to maintain the Earth’s radiation budget and the atmospheric energy transport to define the ocean heat transport as a residual. In its original implementation, the residual method could only be applied to estimate zonally averaged ocean heat transport into the polar cap north of any given latitude because atmosphere energy transport could only be determined for a complete zonal integral. In a recent development, the atmospheric energy transport divergence is estimated on a grid point basis from a globally consistent model analysis that assimilates atmospheric observations and
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radiation variables. From such analysis, the surface energy flux can be estimated at each grid point as the difference between radiation input at the top of the atmosphere and atmospheric transport divergence. Presently, the radiation input can only be accurately estimated for the intensive period of Earth Radiation Budget Experiment from 1985 to 1989. Imposing constraints that the annual averaged surface flux over land should be zero and that the net global air–sea flux should be zero leads to realistic charts of air–sea heat exchange from which ocean heat transport divergence can be estimated. Careful comparison of such ocean heat transport convergence with existing air–sea flux climatologies and with direct estimates of ocean heat transport convergence has not yet been done. There are many outstanding questions on how ocean heat transport will change under changing climate conditions. As atmospheric CO2 has increased, the ocean has warmed up by 14 1022 J over the past 40 years. Such warming represents a heat flux of only 0.1 PW or 0.3 W m 2 averaged over the ocean surface area, so it is unlikely to be detectable in local estimates of air–sea exchange that have uncertainties of 30 W m 2 or in direct estimates of ocean heat transport convergence with uncertainties of 0.3 PW. Instead, local estimates of ocean heat content change over decadal timescales provide a sensitive estimate of how the difference between ocean heat transport convergence and air– sea exchange is changing in a changing climate. There may be changes in ocean circulation that will lead to measurable changes in air–sea heat exchange and ocean heat transport. For example, most coupled climate models predict that the Atlantic meridional overturning circulation will slow down by order of 50% over the next century as atmospheric CO2 increases. Because Atlantic heat transport is presently closely related to the strength of the meridional overturning circulation, Atlantic ocean heat transport could reduce measurably. In addition, the absence of a meridional overturning circulation in coupled climate models leads to much colder (10 1C lower) temperatures in the northern Atlantic that greatly reduce the amount of heat given up by the ocean to the atmosphere in northern latitudes. In fact, there has been a recent suggestion that the Atlantic meridional circulation decreased by 30% since 1992. The heat transport decreased by only about 15% from 1.3 to 1.1 PW, however, as the horizontal gyre circulation increased to transport more heat northward. Thus, the Atlantic heat transport may not reduce proportionately with a decreased meridional overturning circulation, because in the absence of a meridional overturning circulation it is possible that the Atlantic will become more like the Pacific with a horizontal
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gyre circulation that still transports a substantial amount of heat northward. If Atlantic ocean heat transport reduces under changing climate, will the atmospheric energy transport act to compensate with larger northward energy transport? Presently, the Atlantic Ocean circulation transports more than 20% of the maximum energy transport required to balance the Earth’s radiation budget. The hypothesized Bjerkenes compensation mechanism suggests that a reduction in ocean heat transport will be compensated by increased atmospheric energy transport. For extratropical latitudes, atmospheric transport is primarily effected by eddies, cyclones, and anticyclones. Will a reduction in ocean heat transport then be accompanied by increased mid-latitude storminess and increased atmospheric energy transport? Or will the overall radiation budget for the Earth system be fundamentally altered? Clearly, it is of interest to monitor the changes in ocean circulation and heat transport, most importantly in the Atlantic where there are concerns that increasing atmospheric CO2 may lead relatively quickly to substantial changes in ocean circulation and heat transport. A program to monitor the Atlantic meridional overturning circulation and associated heat transport started in 2004 and such monitoring may provide the first evidence for changes in Atlantic ocean heat transport.
See also Agulhas Current. Ekman Transport and Pumping. Florida Current, Gulf Stream and Labrador Current. Indonesian Throughflow. Kuroshio and Oyashio Currents. Meddies and Sub-Surface Eddies. Mesoscale Eddies. Neutral Surfaces and the Equation of State.
Further Reading Bryden HL (1993) Ocean heat transport across 241 N latitude. In: McBean GA and Hantel M (eds.)
Geophysical Monograph Series, Vol. 75 Interactions between Global Climate Subsystems: The Legacy of Hann, pp. 65--75. Washington, DC: American Geophysical Union. Bryden HL and Beal LM (2001) Role of the Agulhas Current in Indian Ocean circulation and associated heat and freshwater fluxes. Deep-Sea Research I 48: 1821--1845. Bryden HL and Imawaki S (2001) Ocean heat transport. In: Siedler G, Church J, and Gould J (eds.) Ocean Circulation and Climate, pp. 455--474. New York: Academic Press. Bryden HL, Longworth HR, and Cunningham SA (2005) Slowing of the Atlantic meridional overturning circulation at 251 N. Nature 438: 655--657. Cunningham SA, Kanzow T, Rayner D, et al. (2007) Temporal variability of the Atlantic meridional overturning circulation at 26.51 N. Science 317: 935--938. Ganachaud A and Wunsch C (2000) Improved estimates of global ocean circulation, heat transport and mixing from hydrographic data. Nature 408: 453--457. Josey SA, Kent EC, and Taylor PK (1999) New insights into the ocean heat budget closure problem from analysis of the SOC air–sea flux climatology. Journal of Climate 12: 2856--2880. Lavı´n A, Bryden HL, and Parrilla G (1998) Meridional transport and heat flux variations in the subtropical North Atlantic. Global Atmosphere and Ocean System 6: 269--293. Levitus S, Antonov J, and Boyer T (2005) Warming of the World Ocean, 1955–2003. Geophysical Research Letters 32 (doi:10.1029/2004GL021592). Shaffrey L and Sutton R (2006) Bjerknes compensation and the decadal variability of the energy transports in a coupled climate model. Journal of Climate 19(7): 1167--1181. Trenberth KE and Caron JM (2001) Estimates of meridional atmosphere and ocean heat transports. Journal of Climate 14: 3433--3443. Trenberth KE, Caron JM, and Stepanaik DP (2001) The atmospheric energy budget and implications for surface fluxes and ocean heat transports. Climate Dynamics 17: 259--276. Vellinga M and Wood RA (2002) Global climatic impacts of a collapse of the Atlantic thermohaline circulation. Climatic Change 54: 251--267.
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HISTORY OF OCEAN SCIENCES H. M. Rozwadowski, Georgia Institute of Technology, Atlanta, Georgia, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1206–1210, & 2001, Elsevier Ltd.
Oceanography is the scientific study of the ocean, its inhabitants, and its physical and chemical conditions. Since its emergence as a recognized scientific discipline, oceanography has been characterized less by the intellectual cohesiveness of a traditional academic discipline than by multidisciplinary, often large-scale, investigation of a complex and forbidding environment. The term ‘oceanography’ was not applied until the 1880s, at which point it still competed with such alternatives as ‘thassalography’ and ‘oceanology’. In many countries, ‘oceanography’ now encompasses both biological and physical traditions, but this meaning is not universal. In Russia, for instance, it does not include biological sciences; the umbrella term remains ‘oceanology’. Before scientists studied the ocean as a geographic place with an integrated ecosystem, they addressed individual biological, physical, and chemical questions related to the sea. European expansion promoted the study of marine phenomena, initiating a lasting link between commercial interests and oceanography. Institutional interest in the sea by the Royal Society in London flourished briefly from 1660 to 1675, when members, including Robert Hooke and Robert Boyle, discussed marine research, developed instruments to make observations at sea, and conducted experiments on sea water to discover its physical and chemical properties. Sailors and travelers continued to make scattered observations at sea, working from the research plan and equipment established by the Society. In addition, tidal studies were prosecuted consistently from the late seventeenth century to well into the nineteenth. Following decades of quiescence, a renewal of interest in marine phenomena occurred in the mideighteenth century. As part of that century’s growth of astronomy, geophysics, chemistry, geology, and meteorology, investigators began making observations at sea as part of their scientific pursuit of other fields. Initially, observers were traveling gentlemen, naturalists, or eclipse-expedition astronomers; later they were scientific explorers. Beginning with the voyages of Captain James Cook in the last third of the century, British exploration became
characterized by attention to scientific observations. The amount of energy devoted to marine science depended on the level of interest of expedition leaders or individual members, but, in the last quarter of the century, the volume of experiments and observations made at sea increased. Virtually all ocean investigation during this time focused on temperature and salinity of water, reflecting the growth of chemical sciences. The concept of oceanic circulation sustained by differences in density, which was widely discussed in the early nineteenth century, emerged at this time. The study of waves became more important while marine natural history remained at a modest level. Early in the nineteenth century the emphasis on physical and chemical studies continued. Curtailed exploration due to the American Revolutionary War and the Napoleonic Wars slowed work in marine science until another period of rapid expansion between 1815 and 1830. Work then focused largely on currents and salinity, reflecting the fact that marine science was conducted on Arctic expeditions which searched for sperm whaling grounds and the Northwest Passage. Navigation and the pursuit of whales prompted Arctic explorers to investigate water temperature and pressure at depths. Enthusiastic individuals, most frequently ships’ captains such as William Scorseby, continued to carry on observation programs at sea, but the 1830s saw a decrease of interest in marine science by physical scientists, who turned to rival fields of discovery including meteorology and terrestrial magnetism. At this time, however, British zoologists directed particular attention to marine fauna, embarking on small boats and yachts to collect with modified oyster dredges. Pursuit of new species as well as living relatives of fossilized ones inspired naturalists to reach deeper and deeper into the sea. Beginning in the 1850s, both British and American hydrographic institutions began deep sea sounding experiments which were guided by the promise of submarine telegraphy, although initiated in support of whaling and navigational concerns. The decades after 1840 saw the gradual awareness by scientists, sailors, entrepreneurs, and governments that the ocean’s depths were commercially and intellectually important places to investigate. An unprecedented increase in popular awareness of the sea accompanied this trend. Prefaced by the rage for seaside vacations which was initiated by railroad access to beaches, cultural interest in the ocean was
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manifested by the popularity of marine natural history collecting and the new maritime novels as well as the vogue for yachting and the personal experiences of growing numbers of ocean travelers and emigrants. Submarine telegraphy and public interest in maritime issues helped scientists argue successfully for government funding of oceanographic voyages. Declining fisheries emphasized the need to know more about the biology of the seas. Until 1840, Britain led marine science, although important work was conducted by investigators in other countries, especially Scandinavian countries. In Norway, for example, G. O. Sars carried out important research for fisheries and also studied unusual creatures dredged from very deep water. American marine science began to threaten British dominance after mid-century. Matthew Fontaine Maury was the first investigator to compile wind, current, and whale charts to improve navigation and commerce. He also dispatched the first trans-Atlantic sounding voyages in the early 1850s. Under Alexander Dallas Bache’s tenure, the US Coast Survey began to include, alongside routine charting work, special studies such as detailed surveys of the Gulf Stream, microscopic examinations of bottom sediments, and dredging cruises with the renowned zoologist Louis Agassiz and his son Alexander. Spencer F. Baird, Secretary of the Smithsonian Institution and United States Fish Commissioner, oversaw American efforts to study marine fauna. In Britain, while physical work was undertaken by the Hydrographic Office and Admiralty exploring expeditions, marine biological science centered around the British Association for the Advancement of Science Dredging Committee from 1839 until the mid-1860s. After that time, a Royal Society–Admiralty partnership took the lead and dispatched a series of summer expeditions. These culminated in the famous fouryear circumnavigation of HMS Challenger (1872– 1876), the first expedition sent out with a mandate to study the world’s oceans. The United States, Britain, and other nations as well, quickly followed the example of the Naples Zoological Station and set up coastal marine biological laboratories in the 1880s and 1890s. In the last quarter of the nineteenth century, many nations sponsored oceanographic voyages modeled after that of Challenger. The United States, Russia, Germany, Norway, France, and Italy contributed to the effort to define the limits and contents of the oceans. Late in the nineteenth century, however, the Scandinavian countries promoted a new style of ocean science to replace the great voyage tradition. Mounting concerns about depleted fisheries inspired national efforts in many countries to study the
biology of fish species as well as their migration. Sweden initiated the formation of what became in 1902 the International Council for the Exploration of the Seas (ICES), which coordinated research undertaken by eight northern European member nations. Although not a member, the United States also continued active biological research. Victor Henson’s discovery of plankton and efforts to quantify its distribution and the subsequent realization of how to use physical oceanography to investigate the movements of fish populations gave oceanographers confidence with which to study the ocean as an undivided system. World War I disrupted the international community of oceanographers but provided the impetus for developing echo-sounding technology for submarine detection, which had been pioneered for ice detection partly in response to the Titanic disaster. By the late 1920s, echo-sounders revolutionized the study of underwater topography and helped scientists to recognize the rift valleys of midocean ridges, showing them to be active, unstable regions. Fishermen took up this technology for locating schools of pelagic fish; later scientists adapted echo-sounding to create fishery-independent surveying tools. Although government funding of oceanographic research dropped back almost to pre-war levels, the late 1920s saw the appearance of the first oceanographic institutions, sponsored mostly by foundations and private individuals. In the United States, the Scripps Institution changed its mission from biological research to oceanography in 1925 and, five years later, the Woods Hole Oceanographic Institution was established on the Atlantic. The availability of oceangoing vessels in the 1930s spurred development of American oceanography on a new scale. Fisheries research, which blossomed in Europe in the 1930s, included important theoretical advances which provided the foundation for later fish population dynamics modeling as well as open ocean research such as Sir Alister Hardy’s Continuous Plankton Recorder surveys and Johannes Schmidt’s publicly acclaimed search for the mid-Atlantic spawning ground of the European eel. Economic depression affected practical government science, such as fisheries research, as well as private projects, such as several of Schmidt’s voyages, which were sponsored by the Carlsberg Foundation. As a consequence, oceanographic work, which was particularly expensive, slowed down dramatically until preparations began for World War II. As with other sciences, World War II partnerships helped forge a new relationship between governments and oceanography. Oceanographic work during and after the war carried the imprint of wartime
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HISTORY OF OCEAN SCIENCES
government support and policy in both its problem selection and its scale. Physical studies gained and maintained precedence over biological ones. The areas of inquiry promoted by wartime efforts related to submarine and antisubmarine tactics. New research began on underwater acoustics, and ocean floor sediment charts were compiled from existing data. Wave studies also received precedence for their value in predicting surf conditions for landings. After the war, oceanographers and academic institutions, newly accustomed to generous funding, learned to accept and even encourage government support. The foundation of the National Science Foundation, which became the major federal supporter of marine science in the USA, exemplified the new high levels of support for this technology-intensive science. Oceanography became characterized by large-scale, expensive research projects such as the 1960s deepsea drilling by proponents of the theory of seafloor spreading. Major international projects also became an integral part of ocean sciences, in conjunction with postwar internationalization of science and other sectors. The International Indian Ocean Expedition, for example, targeted the least well-studied of the world’s oceans. By 1950 physical oceanographers were aware of the wanderings of the Gulf Stream and pressed urgently for systematic exploration of ocean variability. Development of deep ocean mooring technology provided the opportunity to study currents and temperatures continuously. This work showed that mesoscale variability was important and was incorporated into general circulation models. Progress in postwar physical oceanography in general proceeded in step with technological improvements, particularly the development of computing power. Another example, that of radioactive tracers and ‘experiments’ of atomic bomb-testing at sea, led to an intensification of research into biogeochemical cycling from the 1950s. Through programs such as the Geochemical Ocean Sections Study (GEOSECS), tracer use permitted estimates of mixing and circulation times by the late 1960s and led to programs such as the World Ocean Circulation Experiment (WOCE). Circulation and diffusion studies took on renewed importance as public concern about pollution rose. The idea that it was necessary to understand how the ocean functioned in order to avoid inadvertently destroying it fanned development of biological and chemical oceanography from the late 1960s onward. Much effort was funneled into baseline surveying and monitoring programs, initially focusing on potential threats to human health through eating poisoned seafood but soon broadening to investigations of biological effects of
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contaminants. The fast growth of mariculture from this period also encouraged these studies. Although many oceanographers became interested in open ocean research and theoretical modeling in the postwar period, some in Europe resuscitated efforts, begun within ICES in the 1930s, to integrate hydrographic knowledge with fisheries biology. This work encompassed the practical attempt to guide fishermen to catch more fish as well as the scientific project of trying to relate recruitment fluctuations to environmental factors. Parallel efforts in the United States bore fruit in the California Cooperative Oceanic Fisheries Investigations (CalCOFI), a cooperative government and academic program which investigated causes for the decline of the California coastal sardine fishery. CalCOFI, though, was an exceptional project; in general through the 1960s little significant intellectual exchange occurred between biological oceanography and fisheries science. Within biological oceanography, the legacy of E. Steeman-Nielsen’s radioactive carbon tracer method led to active work measuring primary productivity, with global estimates by J. H. Ryther and others in the late 1960s. Discovery of high biological diversity in the deep sea in the late 1960s led to work which was considerably enhanced by deep-diving submersibles, whose use dramatically changed our perception of the ocean’s depths. In the area of marine geology and geophysics, the two greatest achievements since World War II have been the development of the theory of plate tectonics and the deciphering of Earth’s paleoclimate record from deep-sea sediments. The first of these was set in motion by the Vine and Matthews hypothesis of 1963, which rocked the scientific community but was accepted with remarkable speed. Research submersibles played a central role in proving the theory of seafloor spreading. The 1979 discovery of hydrothermal vents, made in the process of these geophysical investigations, led to the surprise discovery of chemosynthetic ecosystems. Reconstruction of Earth’s paleoclimates, which became possible due to the large accumulation of data and the availability of deep sea cores from the ocean drilling projects, was fueled by growing scientific concerns about global climate change. This work began during the 1960s with C. David Keeling’s launch of time-series measurements of carbon dioxide, which provided the data documenting the increase of atmospheric carbon dioxide attributable to human activities and prompted inquiry into the magnitude of exchange of carbon dioxide between the atmosphere and oceans. Studies of the ocean’s role in global warming and weather production have grown in importance in recent years.
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Accompanying the rise of the environmental movement, oceanography’s multi-faceted attempt to understand the oceans as integrated biological and physical environments made it a compelling discipline to ecologically and environmentally minded scientists from the 1970s onward. Food web investigations became prominent, including those which used mesoscale enclosures such as one designed and run at Canada’s Pacific Biological Station in Nanaimo. Discovery of the microbial character of the pelagic food web added a new dimension to biological oceanography. New understanding of production, including a distinction between recycled nutrients and new production made by Dugdale and Goering in 1967, provided biological oceanography with the mathematical formalism for rigorous, quantitative modeling of ocean productivity and biogeochemical fluxes. Modeling capability encouraged a link to be forged between physical and biological oceanography, because new concepts required that physical processes of mixing and upwelling be integrated into ecosystem models dealing with new production, fish production, or export of organic material from the surface layer. The international Global Ocean Ecosystem Dynamics program and the recent resurgence of fisheries oceanography exemplify relatively successful efforts to bridge trophic levels, from plankton to marine mammals and sea birds, and thereby study the ecosystem as a whole.
Distribution. Primary Production Primary Production Processes.
Methods.
Further Reading Deacon M (1971) Scientists and the Sea, 1650–1900: A Study of Marine Science. London: Academic Press. Herdman W (1923) Founders of Oceanography and their Work: An Introduction to the Science of the Sea. London: Edward Arnold. Idyll CP (1969) Exploring the Ocean World: A History of Oceanography. New York: Thomas Y. Crowell. Lee AJ (1992) The Directorate of Fisheries Research: Its Origins and Development. London: Ministry of Agriculture, Fisheries, and Food. Mills EL (1989) Biological Oceanography: An Early History, 1870–1960. Ithaca: Cornell University Press. Schlee S (1973) The Edge of an Unfamiliar World: A History of Oceanography. New York: E.P. Dutton. Smith TD (1994) Scaling Fisheries: The Science of Measuring the Effects of Fishing, 1855–1955. Cambridge: Cambridge University Press. Went AEJ (1972) Seventy years agrowing: a history of the International Council for the Exploration of the Sea, 1902–1972. Rapport et Proce`s-Verbaux des Re´unions, Conseil Permanent International pour l’Exploration de la Mer 165: 1--252. National Research Council 50 (2000) Years of Ocean Discovery: National Science Foundation, 1950–2000. Washington, DC: National Academy Press, 2000.
See also Continuous Plankton Recorders. Fisheries and Climate. Indonesian Throughflow. International Organizations. Law of the Sea. Leeuwin Current. Maritime Archaeology. Primary Production
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HOLOCENE CLIMATE VARIABILITY M. Maslin, C. Stickley, and V. Ettwein, University College London, London, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1210–1217, & 2001, Elsevier Ltd.
Introduction Until a few decades ago it was generally thought that significant large-scale global and regional climate changes occurred at a gradual pace within a timescale of many centuries or millennia. Climate change was assumed to be scarcely perceptible during a human lifetime. The tendency for climate to change abruptly has been one of the most surprising outcomes of the study of Earth history. In particular, paleoceanographic records demonstrate that our present interglacial, the Holocene (the last B10 000 years), has not been as climatically stable as first thought. It has been suggested that Holocene climate is dominated by millennial-scale variability, with some authorities suggesting that this is a 1500 year cyclicity. These pronounced Holocene climate changes can occur extremely rapidly, within a few centuries or even within a few decades, and involve regional-scale changes in mean annual temperature of several degrees Celsius. In addition, many of these Holocene climate changes are stepwise in nature and may be due to thresholds in the climate system. Holocene decadal-scale transitions would presumably have been quite noticeable to ancient civilizations. For instance, the emergence of crop agriculture in the Middle East corresponds very closely with a sudden warming event marking the beginning of the Holocene, and the widespread collapse of the first urban civilizations, such as the Old Kingdom in Egypt and the Akkadian Empire, coincided with a cooling event at around 4300 BP. In addition, paleo-records from the late Holocene demonstrate the possible influence of climate change on the collapse of the Mayan civilization (Classic Period), while Andean ice core records suggests that alternating wet and dry periods influenced the rise and fall of coastal and highland cultures of Ecuador and Peru. It would be foolhardy not to bear in mind such sudden stepwise climate transitions when considering the effects that humans might have upon the present climate system, via the rapid generation of greenhouse gases for instance. Judging by what we
have already learnt from Holocene records, it is not improbable that the system may gradually build up over hundreds of years to a ‘breaking point’ or threshold, after which some dramatic change in the system occurs over just a decade or two. At the threshold point, the climate system is in a delicate and somewhat critical state. It may take only a relatively minor ‘adjustment’ to trigger the transition and tip the system into abrupt change. This article summarizes the current paleoceanographic records of Holocene climate variability and the current theories for their causes. Concentrating on records that cover a significant portion of the Holocene. The discussion is limited to centennial– millennial-scale variations. Figure 1 illustrates the Holocene and its climate variability in context of the major global climatic changes that have occurred during the last 2.5 million years. Short-term variations such as the North Atlantic Oscillation and the El Nin˜o-Southern Oscillation will not be discussed.
The Importance of the Oceans and Holocene Paleoceanography Climate is created from the effects of differential latitudinal solar heating. Energy is constantly transferred from the equator (relatively hot) toward the poles (relatively cold). There are two transporters of such energy – the atmosphere and the oceans. The atmosphere responds to an internal or external change in a matter of days, months, or may be a few years. The oceans, however, have a longer response time. The surface ocean can change over months to a few years, but the deep ocean takes decades to centuries. From a physical point of view, in terms of volume, heat capacity and inertia, the deep ocean is the only viable contender for driving and sustaining long-term climate change on centennial to millennial timescales. Since the process of oceanic heat transfer largely regulates climate change on longer time-scales and historic records are too short to provide any record of the ocean system prior to human intervention, we turn to marine sediment archives to provide information about ocean-driven climate change. Such archives can often provide a continuous record on a variety of timescales. They are the primary means for the study and reconstruction of the stability and natural variability of the ocean system prior to anthropogenic influences.
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Future 0
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Figure 1 Log time scale cartoon, illustrating the most important climate events in the Quaternary Period. (a, ka, Ma, refer to years ago.) (1) Onset of Northern Hemisphere Glaciation (3.2–2.5 Ma), ushering in the strong glacial–interglacial cycles which are characteristic of the Quaternary Period. (2) Mid-Pleistocene Revolution when the dominant periodicity of glacial–interglacial cycles switched from 41 000 y, to every 100 000 y. The external forcing of the climate did not change; thus, the internal climate feedback’s must have altered. (3) The two closest analogues to the present climate are the interglacial periods at 420 000 to 390 000 years ago (oxygen isotope stage 11) and 130 000 to 115 000 years ago (oxygen isotope stage 5e, also known as the Eemian). (4) Heinrich events and Dansgaard–Oeschger cycles (see text). (5) Deglaciation and the Younger Dryas events. (6) Holocene Dansgaard– Oeschger cycles (see text). (7) Little Ice Age (AD 1700), the most recent climate event that seems to have occurred throughout the Northern Hemisphere. (8) Anthropogenic global warming.
One advantage of marine sediments is that they can provide long, continuous records of Holocene climate at annual (sometimes intraannual) to centennial time-resolutions. However, there is commonly a trade-off between temporal and spatial resolution. Deep-ocean sediments usually represent a large spatial area, but sedimentation rates in the deep-ocean are on average between 0.002 and 0.005 cm y 1, with very productive areas producing a maximum of 0.02 cm y 1. This limits the temporal resolution to a maximum of 200 years per cm (50 y cm 1 for productive areas). Mixing by the process of bioturbation will reduce the resolution further. On continental shelves and in bays and other specialized sediment traps such as anoxic basins and fiords, sedimentation rates can exceed 1 cm y 1 providing temporal resolution of over 1 y cm 1. More local conditions are recorded in laminated marine sediments formed in anoxic environments, where biological activity can not disturb the sediments. For example, Pike and Kemp (1997) analysed annual and intraannual variability within the Gulf of California from laminated sediments containing a record of diatom-mat accumulation. Time series analysis highlighted a decadal-scale variability in mat-deposition associated with Pacific-wide changes in surface water circulation, suggested to be influenced by solar-cycles. In addition, anoxic sediments from the Mediterranean Ridge (ODP Site 971) reveal
seasonal-scale variability during the late Quaternary from a laminated diatom-ooze sapropel. Pearce et al., (1998) inferred changes in the monsoon-related nutrient input to the Mediterranean Basin via the Nile River as the main cause of the variations in the laminated sediments, which suggests a wide influence of changes in seasonality. Other potentially extremely high-resolution studies will come from Saanich Inlet, a Canadian fiord and Prydz Bay in Antarctica, sites recently drilled by the Ocean Drilling Program. However, the main drawback to such high-resolution locations is that they contain highly localized environmental and climate information. An additional problem associated mainly with continental margins is reworking, erosion, and redistribution of the sediment by mass density flows such as turbidities and slumps. Hence we concentrate on wider-scale records of Holocene climate change.
Holocene Climatic Variability Initial studies of the Greenland ice core records concluded the absence of major climate variation within the Holocene. This view is being progressively eroded, particular in the light of new information being obtained from marine sediments (Figure 2). Long-term trends indicate an early to mid-Holocene climatic optimum with a cooling trend in the late
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Northwest African Climate
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(F)
Figure 2 Comparison of summer insolation for 651N with north-west African climate (deMenocal et al., 2000) and North Atlantic climate (V29-191, Bond et al., 1997; NEAP 15 K, Bianchi and McCave, 1999; GISP2, O’Brien et al., 1996). Note the similarity of events labeled 1 to 8 and the Little Ice Age (LIA).
Holocene. Superimposed on this trend are several distinct oscillations or climatic cooling steps that appear to be of widespread significance (see Figure 2), the most dramatic of which occurred 8200, 5500, and 4400 years ago and between AD 1200 and AD 1650. The event 8200 years ago is the most striking and abrupt, leading to widespread cool and dry conditions lasting perhaps 200 years, before a rapid return to climates warmer and generally moister than at present. This event is noticeably present in the GISP2 Greenland ice cores, from which it appears to have been about half as severe as the Younger Dryas to Holocene transition. Marine records of North African to Southern Asian climate suggest more arid conditions involving a failure of the summer monsoon rains. Cold and/or arid conditions also seem to have occurred in northernmost South America, eastern North America and parts of north-west Europe. In the middle Holocene approximately 5500–5300 years ago there was a sudden and widespread shift in precipitation, causing many regions to become either noticeably drier or moister. The dust and sea surface temperature records off north-west Africa show that the African Humid Period, when much of subtropical West Africa was vegetated, lasted from 14 800 to 5500 years ago and was followed by a 300year transition to much drier conditions (de Menocal et al., 2000). This shift also corresponds to the decline of the elm (Ulmus) in Europe about 5700, and of hemlock (Tsuga) in North America about 5300
years ago. Both vegetation changes were initially attributed to specific pathogen attacks, but it is now thought they may have been related to climate deterioration. The step to colder and drier conditions in the middle of an interglacial period is analogous to a similar change that is observed in records of the last interglacial period referred to as Marine Oxygen Isotope Stage 5e (Eemian). There is also evidence for a strong cold and arid event occurring about 4400 years ago across the North Atlantic, northern Africa, and southern Asia. This cold, and arid event coincides with the collapse of a large number of major urban civilizations, including the Old Kingdom in Egypt, the Akkadian Empire in Mesopotamia, the Early Bronze Age societies of Anatolia, Greece, Israel, the Indus Valley civilization in India, the Hilmand civilization in Afganistan, and the Hongshan culture of China.
Little Ice Age (LIA) The most recent Holocene cold event is the Little Ice Age (see Figures 2 and 3). This event really consists of two cold periods, the first of which followed the Medieval Warm Period (MWP) that ended B1000 years ago. This first cold period is often referred to as the Medieval Cold Period (MCP) or LIAb. The MCP played a role in extinguishing Norse colonies on Greenland and caused famine and mass migration in Europe. It started gradually before AD 1200 and ended at about AD 1650. This second cold period,
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Greenland temperature
Bermuda Rise SST
(˚C)
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_ 31
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_6
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Figure 3 Comparison of Greenland temperatures, the Bermuda Rise sea surface temperatures (SST) (Keigwin, 1996), and west African and a sea surface temperature (deMenocal et al., 2000) for the last 2500 years. LIALittle Ice Age; MWPMedieval Warm Period. Solid triangles indicate radiocarbon dates.
may have been the most rapid and the largest change in the North Atlantic during the Holocene, as suggested from ice-core and deep-sea sediment records. The Little Ice Age events are characterized by a drop in temperature of 0.5–11C in Greenland and a sea surface temperature falls of 41C off the coast of west Africa and 21C off the Bermuda Rise (see Figure 3).
Holocene Dansgaard–Oeschger Cycles The above events are now regarded as part of the millennial-scale quasiperiodic climate changes characteristic of the Holocene (see Figure 2) and are thought to be similar to glacial Dansgaard–Oeschger (D/O) cycles. The periodicity of these Holocene D/O cycles is a subject of much debate. Initial analysis of the GISP2 Greenland ice core and North Atlantic sediment records revealed cycles at approximately the same 1500 (7500)-year rhythm as that found within the last glacial period. Subsequent analyses have also found a strong 1000-year cycle and a 550year cycle. These shorter cycles have also been recorded in the residual d14 C data derived from dendrochronologically calibrated bidecadal tree-ring measurements spanning the last 11 500 years. In general, during the coldest point of each of the millennial-scale cycles shown in Figure 2, surface water temperatures of the North Atlantic were about 2– 41C cooler than during the warmest part. One cautionary note is that Wunsch has suggested a more radical explanation for the pervasive 1500year cycle seen in both deep-sea and ice core, glacial and interglacial records. Wunsch suggests that the
extremely narrow spectral lines (less than two bandwidths) that have been found at about 1500 years in many paleo-records may be due to aliasing. The 1500-year peak appears precisely at the period predicted for a simple alias of the seasonal cycle sampled inadequately (under the Nyquist criterion) at integer multiples of the common year. When Wunsch removes this peak from the Greenland ice core data and deep-sea spectral records, the climate variability appears as expected to be a continuum process in the millennial band. This work suggests that finding a cyclicity of 1500 years in a dataset may not represent the true periodicity of the millennialscale events. The Holocene Dansgaard–Oeschger events are quasi periodic, with different and possibly stochastic influences.
Causes of Millennial Climate Fluctuation during the Holocene As we have already suggested, deep water circulation plays a key role in the regulation of global climate. In the North Atlantic, the north-east-trending Gulf Stream carries warm and relatively salty surface water from the Gulf of Mexico up to the Nordic seas. Upon reaching this region, the surface water has cooled sufficiently that it becomes dense enough to sink, forming the North Atlantic Deep Water (NADW). The ‘pull’ exerted by this dense sinking maintains the strength of the warm Gulf Stream, ensuring a current of warm tropical water into the North Atlantic that sends mild air masses across to the European continent. Formation of the NADW can be weakened by two processes. (1) The presence
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HOLOCENE CLIMATE VARIABILITY
of huge ice sheets over North America and Europe changes the position of the atmospheric polar front, preventing the Gulf Stream from traveling so far north. This reduces the amount of cooling and the capacity of the surface water to sink. Such a reduction of formation occurred during the last glacial period. (2) The input of fresh water forms a lens of less-dense water, preventing sinking. If NADW formation is reduced, the weakening of the warm Gulf Stream causes colder conditions within the entire North Atlantic region and has a major impact on global climate. Bianchi and McCave, using deep-sea sediments from the North Atlantic, have shown that during the Holocene there have been regular reductions in the intensity of NADW (Figure 2E), which they link to the 1500-year D/O cycles identified by O’Brien and by Bond (1997). There are two possible causes for the millennial-scale changes observed in the intensity of the NADW: (1) instability in the North Atlantic region caused by varying
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freshwater input into the surface waters; and (2) the ‘bipolar seesaw’. There are a number of possible reasons for the instability in the North Atlantic region caused by varying fresh water input into the surface waters:
• • • •
Internal instability of the Greenland ice sheet, causing increased meltwater in the Nordic Seas that reduces deep water formation. Cyclic changes in sea ice formation forced by solar variations. Increased precipitation in the Nordic Seas due to more northerly penetration of North Atlantic storm tracks. Changes in surface currents, allowing a larger import of fresher water from the Pacific, possibly due to reduction in sea ice in the Arctic Ocean.
The other possible cause for the millennial-scale changes is an extension of the suggested glacial intrinsic millennial-scale ‘bipolar seesaw’ to the
0.8
Northward heat flux (1015 W)
0.4 N. Atlantic heat piracy
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(2) CLIMAP LGM (3) Cold tropics LGM (4) Heinrich event (HL)
_ 0.4
(5) Heinrich event (HB) (6) D/O event
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40˚S Latitude
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Figure 4 Atlantic Ocean poleward heat transport (positive indicates a northward movement) as given by the ocean circulation model (Seidov and Maslin, 1999) for the following scenarios: (1) present-day (warm interglacial) climate; (2) last glacial maximum (LGM) with generic CLIMAP data; (3) ‘Cold tropics’ LGM scenario; (4) a Heinrich-type event driven by the meltwater delivered by icebergs from decaying Laurentide ice sheet; (5) a Heinrich-type event driven by meltwater delivered by icebergs from decaying Barents Shelf ice sheet or Scandinavian ice sheet; (6) a general Holocene or glacial Dansgaard–Oeschger (D/O) meltwater confined to the Nordic Seas. Note that the total meridional heat transport can only be correctly mathematically computed in the cases of cyclic boundary conditions (as in Drake Passage for the global ocean) or between meridional boundaries, as in the Atlantic Ocean to the north of the tip of Africa. Therefore the northward heat transport in the Atlantic ocean is shown to the north of 301S only.
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Holocene. One of the most important finds in the study of glacial millennial-scale events is the apparent out-of-phase climate response of the two hemispheres seen in the ice core climate records from Greenland and Antarctica. It has been suggested that this bipolar seesaw can be explained by variations in the relative amount of deep water formation in the two hemispheres and heat piracy (Figure 4). This mechanism of altering dominance of the NADW and the Antarctic Bottom Water (AABW) can also be applied to the Holocene. The important difference with this theory is that the trigger for a sudden ‘switching off’ or a strong decrease in rate of deep water formation could occur either in the North Atlantic or in the Southern Ocean. AABW forms in a different way than NADW, in two general areas around the Antarctic continent: (1) near-shore at the shelf–ice, sea-ice interface and (2) in open ocean areas. In near-shore areas, coastal polynya are formed where katabatic winds push sea ice away from the shelf edge, creating further opportunity for sea ice formation. As ice forms, the surface water becomes saltier (owing to salt rejection by the ice) and colder (owing to loss of heat via latent heat of freezing). This density instability causes sinking of surface waters to form AABW, the coldest and
saltiest water in the world. AABW can also form in open-ocean Antarctic waters; particularly in the Weddell and Ross seas; AABW flows around Antarctica and penetrates the North Atlantic, flowing under the less dense NADW. It also flows into the Indian and Pacific Oceans, but the most significant gateway to deep ocean flow is in the south-west Pacific, where 40% of the world’s deep water enters the Pacific. Interestingly, Seidov and colleagues have shown that the Southern Ocean is twice as sensitive to meltwater input as is the North Atlantic, and that the Southern Ocean can no longer be seen as a passive player in global climate change. The bipolar seesaw model may also be self sustaining, with meltwater events in either hemisphere, triggering a train of climate changes that causes a meltwater event in the opposite hemisphere, thus switching the direction of heat piracy (Figure 5).
Conclusion The Holocene, or the last 10 000 years, was once thought to be climatically stable. Recent evidence, including that from marine sediments, have altered this view, showing that there are millennial-scale
Internal ice sheet oscillation ?
North Atlantic
Heinrich Events
Reduced NADW
Antarctic Warm Events Antartica melting or surging?
Southern Ocean warms up
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Sea level
Meltwater
Deep water oscillating system
Greenland / Iceland melting or surging?
Warmer N. Atlantic
N. Hemisphere heat piracy
Increased NADW
Reduced AABW
? Internal ice sheet oscillation
Figure 5 Possible deep water oscillatory system explaining the glacial and interglacial Dansgaard–Oeschger cycles. Additional loop demonstrates the possible link between interglacial Dansgaard–Oeschger cycles and Heinrich events.
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HOLOCENE CLIMATE VARIABILITY
Solar cycles North Atlantic Oscillation
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D/O cycles
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? T h e _6
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Figure 6 Spectrum of climate variance showing the climatic cycles for which we have good understanding and the ‘gap’ between hundreds and thousands of years for which we still do not have adequate understanding of the causes.
climate cycles throughout the Holocene. In fact we are still in a period of recovery from the last of these cycles, the Little Ice Age. It is still widely debated whether these cycles are quasiperiodic or have a regular cyclicity of 1500 years. It is also still widely debated whether these Holocene Dansgaard–Oeschger cycles are similar in time and characteristic to those observed during the last glacial period. A number of different theories have been put forward for the causes of these Holocene climate cycles, most suggesting variations in the deep water circulation system. One suggestion is that these cycles are caused by the oscillating relative dominance of North Atlantic Deep Water and Antarctic Bottom Water. Holocene climate variability still has no adequate explanation and falls in the ‘gap’ of our knowledge between Milankovitch forcing of ice ages and rapid variations such as El Nin˜o and the North Atlantic Oscillation (Figure 6). Future research is essential for understanding these climate cycles so that we can better predict the climate response to anthropogenic ‘global warming.’
See also Antarctic Circumpolar Current. Calcium Carbonates. Cenozoic Oceans – Carbon Cycle Models. Current Systems in the Atlantic Ocean. Deep Convection. Deep-Sea Drilling Methodology. Deep-Sea Drilling Results. MillennialScale Climate Variability. Ocean Circulation. Ocean Subduction. Ocean Margin Sediments. Paleoceanography. Sediment Chronologies. Water Types and Water Masses.
Further Reading Adams J, Maslin MA, and Thomas E (1999) Sudden climate transitions during the Quaternary. Progress in Physical Geography 23(1): 1--36. Alley RB and Clark PU (1999) The deglaciation of the Northern hemisphere: A global perspective. Annual Review of Earth and Planetary Science 27: 149--182. Bianchi GG and McCave IN (1999) Holocene periodicity in North Atlantic climate and deep-ocean flow south of Iceland. Nature 397: 515--523. Broecker W (1998) Paleocean circulation during the last deglaciation: a bipolar seesaw? Paleoceanography 13: 119--121. Bond G, Showers W, Cheseby M, et al. (1997) A pervasive millenial-scale cycle in North Atlantic Holocene and glacial climates. Science 278: 1257--1265. Chapman MR and Shackleton NJ (2000) Evidence of 550 year and 1000 years cyclicities in North Atlantic pattern during the Holocene. Holocene 10: 287--291. Cullen HM, et al. (2000) Climate change and the collapse of the Akkadian Empire: evidence from the deep sea. Geology 28: 379--382. Dansgaard W, Johnson SJ, and Clausen HB (1993) Evidence for general instability of past climate from a 250-kyr ice-core record. Nature 364: 218--220. deMenocal P, Ortiz J, Guilderson T, and Sarnthein M (2000) Coherent high- and low-latitude climate variability during the Holocene warm period. Science 288: 2198--2202. Keigwin LD (1996) The Little Ice Age and Medieval warm period in the Sargasso sea. Science 274: 1504--1507. Maslin MA, Seidov D, and Lowe J (2001) Synthesis of the nature and causes of sudden climate transitions during the Quaternary. AGU Monograph: Oceans and Rapid Past and Future Climate Changes: North–South Connections. Washington DC: American Geophysical. Union No. 119.
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O’Brien SR, Mayewski A, and Meeker LD (1996) Complexity of Holocene climate as reconstructed from a Greenland ice core. Science 270: 1962--1964. Pearce RB, Kemp AES, Koizumi I, Pike J, Cramp A, and Rowland SJ (1998) A lamina-scale, SEM-based study of a late quaternary diatom-ooze sapropel from the Mediterranean ridge, Site 971. In: Robertson AHF, Emeis K-C, Richter C, and Camerlenghi A (eds.) Proceedings of the Ocean Drilling Program, Scientific Results 160, 349--363. Peiser BJ (1998) Comparative analysis of late Holocene environmental and social upheaval: evidence for a disaster around 4000 BP. In: Peiser BJ, Palmer T, and Bailey M (eds.) Natural Catastrophes during Bronze Age Civilisations, 117--139.
Pike J and Kemp AES (1997) Early Holocene decadal-scale ocean variability recorded in Gulf of California laminated sediments. Paleoceanography 12: 227--238. Seidov D and Maslin M (1999) North Atlantic Deep Water circulation collapse during the Heinrich events. Geology 27: 23--26. Seidov D, Barron E, Haupt BJ, and Maslin MA (2001) Meltwater and the ocean conveyor: past, present and future of the ocean bi-polar seesaw. AGU Monograph: Oceans and Rapid Past and Future Climate Changes: North–South Connections. Washington DC: American Geophysical. Union No. 119.. Wunsch C (2000) On sharp spectral lines in the climate record and millennial peak. Paleoceanography 15: 417--424.
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HYDROTHERMAL VENT BIOTA R. A. Lutz, Rutgers University, New Brunswick, NJ, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1217–1227, & 2001, Elsevier Ltd.
On 17 February 1977, the deep-submergence vehicle Alvin descended 2500 m to the crest of the Galapagos Rift spreading center to first visit an ecosystem that would forever change our view of life in the deep sea. Cracks and crevices in the ocean floor were emanating fluids with temperatures up to 171C. None of the bizarre organisms clustering around these ‘hydrothermal vents’ had ever been encountered; they comprised new species, genera, families, superfamilies, and bizarre ‘tubeworms,’ up to 2 m long, which were subsequently placed in a new phylum (Vestimentifera) (Figure 1). Since the Galapagos Rift discovery, numerous hydrothermal vent sites have been found throughout the world’s oceans and over 500 new species have been described from these regions. Figure 2 depicts many of the major hydrothermal systems from which organisms have been collected to date. Fluids with temperatures as high as 4031C exit from polymetallic sulfide chimneys in many of these regions (Figure 3). Most ecosystems on earth ultimately rely on photosynthesis, with the energy source being solar. In marked contrast, deep-sea hydrothermal ecosystems
are based predominantly on chemosynthesis, with the energy source being geothermal. Many of the chemosynthetic microbes are fueled by hydrogen sulfide, which is present at low-temperature vents in concentrations up to several hundred micromoles per liter and at high-temperature vents in concentrations up to 100 milimoles per liter. These microbial organisms can be either ‘free-living’ (in the water or on the surface of various substrates) or symbiotic in association with certain vent organisms. The vestimentiferan tubeworms Riftia pachyptila and Tevnia jerichonana (Figures 1, 4, 5, and 6), for example, each have a specialized ‘tissue,’ known as the trophosome, which is comprised entirely of chemosynthetic bacteria. The tubeworms have no mouth, no digestive system, and no anus; in short, no opening to the external environment. Hydrogen sulfide diffuses across cell membranes and is transported via the hemoglobin-containing circulatory system to the trophosome, where it is utilized by the associated symbionts. Mussels (Bathymodiolus thermophilus) (Figures 7 and 8) and vesicomyid clams (Calyptogena magnifica) (Figure 9), common along both the Galapagos Rift and East Pacific Rise (EPR), represent two of the other dominant members of the vent megafauna that house chemosynthetic symbionts. In the case of each of these bivalves, the symbionts are associated with the gills and both species have modified feeding apparatuses relative to those of shallow-water related species (likely a result of their
Figure 1 A cluster of vestimentiferan tubeworms (Riftia pachyptila and Tevnia jerichonana), together with a zoarcid fish (Thermarces andersoni) and brachyuran crabs (Bythograea thermydron) inhabiting low-temperature hydrothermal vents at 91500 N along the East Pacific Rise (depth 2500 m).
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Lucky Strike
Southern Explorer Ridge (SEXP) Juan de Fuca Ridge (JDFR) Gorda Ridge (GORD) Sunrise Okinawa Trough
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21° 13° Northern East 11° Pacific Rise (NEPR) 9° 7° Lau Basin 11° 14° 16° 17° 18° 20° 21° Southern East 22° Pacific Rise (SEPR) 31° 32°
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Figure 2 Deep-sea hydrothermal vent sites along mid-oceanic and back-arc ridge systems from which vent organisms have been collected to date. Numbers indicate approximate latitude of site along the East Pacific Rise.
predominant reliance on the associated symbionts for nutrition). Closely related mussels and clams within the families Mytilidae and Vesicomyidae are common constituents of the fauna associated with vents along mid-oceanic ridge and back-arc spreading centers (as well as at many cold-water hydrocarbon seeps) throughout the world’s oceans. All of
Figure 3 A ‘black smoker’ polymetallic sulfide chimney at 201500 N along the East Pacific Rise (depth 2615 m). Temperatures as high as 4031C have been recorded at the orifice of such edifices from which mineral-rich fluids emanate violently in many deep-sea hydrothermal systems.
these bivalve mollusks retrieved to date appear to contain thiotrophic (‘sulfur-feeding’) or methanotrophic (‘methane-feeding’) chemosynthetic symbionts. It should be mentioned that, while chemosynthetic bacteria play a critical role in food chains associated with many vent systems, there are numerous other microbes at the vents belonging to two other recognized kingdoms. Both eukaryotes and Archaea occupy specialized niches within various vent ecosystems and certain Archaea, which have been isolated from environments associated with hightemperature sulfide chimneys, have been reported to occupy the most primitive ‘node’ on the phylogenetic tree of life. Such reports have led to considerable speculation as to whether or not life itself may have originated in hydrothermal vent environments. At various high-temperature vent sites, numerous organisms colonize the sides of active sulfide edifices. Figure 4 depicts a sulfide edifice (named ‘Tubeworm Pillar’) that is 11 m high, the sides of which are covered with tubeworms (both R. pachyptila and T. jerichonana), crabs, zoarcid fish, and a spectrum of smaller, associated vent fauna. At other sites, the sides of ‘black smoker’ chimneys are frequently covered with the tubes of polychaetes, such as Alvinella pompejana, a ‘Pompeii’ worm named after the
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Figure 4 A portion of an 11 m high polymetallic sulfide edifice known as Tubeworm Pillar at 91500 N along the East Pacific Rise (depth 2500 m). The top two-thirds of the edifice is covered with vestimentiferan tubeworms, both Riftia pachyptila (larger organisms) and Tevnia jerichonana (smaller organisms), as well as numerous brachyuran crabs (Bythograea thermydron) and zoarcid fish (Thermarces andersoni).
submersible Alvin (Figure 10). This organism has been reported to routinely withstand long-term exposure to temperatures ranging from 21C to 351C and short-term exposure to temperatures in excess of 1001C, rendering this annelid perhaps the most eurythermal organism on the planet. In addition to the numerous species of bivalves, such as mussels and clams mentioned earlier, there are myriad other common mollusks at vents, including numerous gastropods, particularly archeogastropod limpets. Thirteen different gastropod species have been reported from vents along the narrow range of the East Pacific Rise from 91170 N to
91540 N. These organisms can achieve high population densities and are found on a wide variety of substrates including the tubes of Riftia pachyptila and the shells of Bathymodiolus thermophilus (Figures 8 and 11). Limpets graze on free-living microbes that coat the majority of surfaces associated with both low-temperature and high-temperature (e.g., sulfide chimney) vents. Over 100 species of mollusks have been collected to date from the mid-oceanic and back-arc spreading centers visited to date. Virtually all of the invertebrates inhabiting deepsea hydrothermal vents have planktonic larval stages. These free-swimming stages serve as the primary
Figure 5 Higher magnification of the side of Tubeworm Pillar (depicted in Figure 4), showing tubeworms (Riftia pachyptila and Tevnia jerichonana) and scavenging brachyuran crabs (Bythograea thermydron).
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Figure 6 Close-up image of a cluster of Tevnia jerichonana, together with a brachyuran crab (Bythograea thermydron) and a zoarcid fish (Thermarces andersoni). Note the ‘accordion-like’ morphology of the tube of this species of vestimentiferan tubeworm.
Figure 7 A dense population of mussels (Bathymodiolus thermophilus) inhabiting a low-temperature hydrothermal vent field along the East Pacific Rise. Associated fauna in the field of view include tubeworms (Riftia pachyptila), brachyuran crabs (Bythograea thermydron), zoarcid fish (Thermarces andersoni), and a galatheid crab (Munidopsis subsquamosa) (lower left).
means of dispersal between isolated vent systems, which can be separated by hundreds of kilometers. The larval stages can be either planktotrophic (feeding within the water column) or lecithotrophic (utilizing yolk reserves for nutrition). In either case, it appears that the early life history stages of the vast majority of vent organisms are capable of staying in the water column for considerable lengths of time (likely months in the case of many vent species). Such extended planktonic durations facilitate passive transport over vast distances via ocean currents (velocities of ocean currents between 15 and 30 cm s1 are commonly encountered along the crest of ridge systems). Numerous crustaceans inhabit the vent environment and represent perhaps the dominant scavengers
Figure 8 Close-up of mussels (Bathymodiolus thermophilus) attached to the tubes of the tubeworm Riftia pachyptila. Limpets (Lepetodrilus elevatus) are seen attached to the external surfaces of both the mussel shells and tubeworm tubes.
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Figure 9 Vesicomyid clams (Calyptogena magnifica) line cracks and crevices from which low-temperature fluids are venting in an area known as Clam Acres at 201500 N along the East Pacific Rise (depth 2615 m). This species is common at many vent sites that are believed to be in relatively late stages of succession along the Galapagos Rift and the entire stretch of the East Pacific Rise from 211N to 181S.
of the ecosystem. Brachyuran crabs (e.g., Bythograea thermydron on the Galapagos Rift and the East Pacific Rise) (Figures 1, 5, 6, and 7) appear to be one of the earliest colonizers of hydrothermal vent environments. When new vents were formed during the April, 1991 volcanic eruption along the crest of the East Pacific Rise at 91500 N, several regions were referred to as ‘crab nurseries’ owing to the relatively large abundance of crab larvae (megalopae) in areas
of low-temperature discharge. Eleven months after the eruptive event, the region was populated by tremendous numbers of large crabs. These were frequently observed to be holding with their claws various substrates, such as empty tubeworm tubes or pieces of basalt covered with microbial mats, and voraciously scraping the surfaces with their mouth parts, suggesting that microbes and/or their products may represent an important source of nutrition for
Figure 10 The polychaete Alvinella pompejana emerging from its tube on the side of a black smoker chimney at 91500 N along the East Pacific Rise. It has been suggested that the organism may be the most eurythermal invertebrate on the planet, capable of withstanding short-term exposures to temperatures ranging from less than 21C to in excess of 1001C.
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Figure 11 Limpets (Lepetodrilus elevatus) coating the surfaces of the tubes of Riftia pachyptila and the shell of Bathymodiolus thermophilus (left). A zoarcid fish (Thermarces andersoni) is seen emerging from the tubeworm tubes.
these organisms. They are also frequently seen extending their claws into the tubes of Riftia pachyptila and will readily devour the tissues of any tubeworm or mussel damaged during routine submersible activities. They are easily captured in ‘crab traps’ baited with a wide variety of dead fish. Intraspecific attacks appear to be relatively common occurrences and crabs with only one claw or a missing leg are frequently seen on the bottom, presumably reflecting an encounter with another crab. Galatheid crabs (e.g., Munidopsis subsquamosa) (Figure 12) inhabit peripheral areas of the vents during earlier stages of succession and, as hydrogen sulfide levels gradually decrease over time, they are
commonly encountered in central areas as well, often among tubeworms and mussels. While these ‘squat lobsters’ may also be scavengers, they are seldom, if ever, caught in traps. One chance encounter with a relatively large dead octopod in the peripheral area of a vent revealed a large quantity of galatheid crabs on the carcass and a noticeable absence of brachyuran crabs or other vent organisms (Figure 13). Close-up imagery revealed that the crabs were actively feeding on the dead tissues of the carcass. Galatheids that have been collected to date have generally been ‘transported’ back to the surface in random regions of the submersible itself. The carapaces of this species are frequently covered with fine, filamentous bacteria (Figure 12).
Figure 12 A galatheid crab (Munidopsis subsquamosa) with filamentous bacteria on its carapace. This species is a common inhabitant of peripheral vent environments and is also frequently observed in ambient deep-sea environments.
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HYDROTHERMAL VENT BIOTA
Figure 13 Numerous galatheid crabs (Munidopsis subsquamosa) on and in the immediate vicinity of the carcass of a cirrate octopod approximately 50 m from an active hydrothermal vent at 91500 N along the East Pacific Rise. Closeup imagery revealed that the crabs were actively feeding on the dead tissues of the carcass.
Numerous species of shrimp (Figure 14) have been encountered at vent sites visited to date along ridge axes in all ocean basins. At several sites (e.g., the TAG hydrothermal vent field) along the Mid-Atlantic Ridge, thousands of shrimp are frequently seen ‘swarming’ on top of one another, completely carpeting the sides of large sulfide chimneys. It has been suggested that a large ‘eye spot’ on the back of the carapace of the shrimp is capable of sensing longwavelength light emitted by high-temperature smokers. Shrimp are also relatively common inhabitants
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of vents visited to date in the western Pacific (e.g., at the Alice Springs vent field along the back-arc spreading center in the Mariana Trough). Other smaller, common vent-endemic crustaceans include numerous species of amphipods, copepods, and leptostracans. Halice hesmonectes, an amphipod common at various vent sites along the East Pacific Rise is frequently seen ‘swarming’ above mussel and tubeworm colonies in regions of active low-temperature venting. It has been reported that these dense swarms represent the highest concentration of planktonic invertebrates in the ocean. Serpulid polychaetes (Figure 15) are common inhabitants of the peripheral area of many vent fields. Commonly referred to as ‘feather dusters,’ they were seen extending over large expanses of lava during the early expeditions to the Galapagos Rift in 1977 and 1979 and were subsequently reported at numerous vents along the East Pacific Rise. Their small tubes, which generally reach lengths of about 5 cm, consist of calcite. When the tentacular plumes are withdrawn into the tube, a small ‘plug’ seals the tube from the external environment. Numerous species of apparently vent-endemic fish have been reported from hydrothermal systems along ridge systems throughout the world’s oceans. Bythites hollisi is a bythitid that was encountered on the first dive to the Galapagos Rift vent field in 1977. Bythitids have been observed in large numbers at several other vent sites, such as the hydrothermal field at 91500 N along the East Pacific Rise (Figure 16). One large ‘pit,’ with a diameter of several meters, from which cloudy, shimmering water was emanating, had a concentration of over 20 bythitids,
Figure 14 The shrimp Alvinocaris lusca perched atop a tube of Riftia pachyptila at a low temperature vent along the East Pacific Rise.
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Figure 15 Serpulid polychaetes in the peripheral area of a vent field at 91500 N along the East Pacific Rise.
which were frequently observed with their heads projecting downward into the cloudiest portions of the water at the base of the pit. Zoarcids (eel pouts) are common members of the vent fauna at various sites along the East Pacific Rise and the Mid-Atlantic Ridge, with several different species having been encountered at the various vent fields visited to date. The common species that inhabits many East Pacific Rise vents is Thermarces andersoni (Figures 1, 4, 6, and 11). From analyses of extensive video footage and stomach contents of retrieved specimens of this species, it appears to commonly feed on a wide range of organisms, including bacteria, amphipods, leptostracans, and shrimp. It has also been observed scavenging on the plumes of specimens of vestimentiferan tubeworms that have
been damaged in the process of sampling or maneuvering with the submersible. Enteroptneusts, commonly referred to as ‘spaghetti worms’ were first observed in 1977 draped over pillow lava in peripheral regions of the Galapagos Rift vent field. Few individuals were observed at any of the other many vent fields visited over the next 20 years throughout the world. In 1997, six years after the volcanic eruption at 91500 N along the East Pacific Rise, numerous enteroptneusts were observed at distances ranging from a few meters to a few hundred meters from active vent sites within the region. No individual organisms were observed living in direct association with venting fluids or at distances in excess of a kilometer from active hydrothermal systems. This unusual, soft-bodied
Figure 16 The vent fish Bythites hollisi emerging from a high-temperature vent region known as Hole-to-Hell at 91500 N along the East Pacific Rise. The species is common at various vent fields along the Galapagos Rift and East Pacific Rise.
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Figure 17 The enteroptneust Saxipendium coronatum (commonly known as a ‘spaghetti worm’) on the surface of basalt about 30 m from an active low-temperature vent field at 91500 N along the East Pacific Rise. The image was taken with a prototype high-resolution video camera system equipped with a macro-lens.
invertebrate may represent an organism that is uniquely adapted to an ecotone between active vents and the ambient deep sea. Analyses of video images of numerous individuals taken with a high-resolution video camera system equipped with a macro-lens (Figure 17) have revealed behavioral patterns suggesting that the organism may be ‘grazing’ directly on basaltic surfaces, potentially consuming microbes or organic substances ultimately originating in the vent environment. One of the most unusual and spectacular organisms inhabiting vent ecosystems is the vestimentiferan tubeworm Riftia pachyptila which thrives at numerous vent fields along the Galapagos Rift and East Pacific Rise (Figures 1, 4, and 8). As mentioned earlier, it lives in a symbiotic relationship with chemosynthetic bacteria concentrated within its body. Such a relationship provides an internal, hydrogen sulfide-nourished ‘garden’ that, in turn, nourishes the tubeworm. Although the mechanism by which the host obtains energy from the bacteria is unclear, the energy transfer appears remarkably efficient. An unique opportunity to determine the growth rate potential of R. pachyptila arose as a result of the April, 1991 volcanic eruptive event at 91500 N along the East Pacific Rise mentioned above. In March,
Figure 18 (Right)Temporal sequence of vent community development at a low-temperature vent in a region known as Hole-to-Hell at 91500 N along the East Pacific Rise. This was the site of a volcanic eruption in April, 1991 that entirely decimated previously existing communities within the region. The field of view of each image and the heading of the camera system utilized are approximately the same for each image taken at the following times: (A) April, 1991; (B) March, 1992; (C) December, 1993; (D) October, 1994; and (E) November, 1995. (Shank et al., 1998.)
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1992, no R. pachyptila were present within the region, which had been devastated by the eruption. By December, 1993, less than 2 years later, huge colonies of this tubeworm had colonized the active lowtemperature vents; the tube lengths of many individuals were in excess of 1.5 m (Figure 18 and 19). Such growth rates of more than 85 cm y1 increase in tube length represent the fastest rates of growth documented for any marine invertebrate. It is interesting to note that one of the other fastest-growing marine invertebrates is the giant clam Tridacna squamosa, a bivalve that has a symbiotic association with photosynthetic algae (zooxanthellae). The efficiency with which energy is transferred from the symbiont to the host in certain invertebrates may well be a contributing factor to the remarkable growth rates of these organisms. The rapid succession of a tubeworm-dominated vent community over the 5-year period following the April, 1991 eruption is dramatically illustrated in Figures 18 and 19. Inhabitants of deep-sea hydrothermal vents are among the most spectacular and unusual organisms on the planet. Given the relatively small number of vent ecosystems found to date, many questions have been raised concerning whether conservation measures need to be taken to protect vent communities from anthropogenic disturbances. Any assessment of the potential consequences of anthropogenic impacts needs to consider that the organisms inhabiting deepsea hydrothermal vents thrive in an environment that is constantly being altered radically by geological process. Periodic devastation of entire biological communities is a relatively common occurrence as volcanic and tectonic processes proceed along active ridge axes. Over the past two decades we have learned that many vent communities are remarkably resilient. Populations of essentially all vent organisms indigenous to regions decimated by massive volcanic eruptions have, like a Phoenix rising from the ashes, reestablished themselves in less than a decade. This remarkable resilience in the face of huge natural disasters has profound implications as one considers the potential impacts of exploitation of
Figure 19 (Left)Temporal sequence of vent community development at a low-temperature vent located approximately 500 m from the community depicted in Figure 18. This was also a region that was buried with fresh lava by the April, 1991 volcanic eruptive event, decimating previously existing communities within the area. The field of view of each image and the heading of the camera system utilized are approximately the same for each image taken at the following times: (A) April, 1991; (B) March, 1992; (C) December, 1993; (D) October, 1994; and (E) November, 1995. (Shank et al., 1998.)
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precious mineral and biological resources associated with active hydrothermal systems throughout the world’s oceans.
Acknowledgments I thank the pilots and crew of the DSV Alvin and the R/V Atlantis for their expertise, assistance, and patience over the years; W. Lange and the Woods Hole Oceanographic Institution for technical expertise and the provision of the camera and recording systems critical to the generation of the majority of images presented in this article; Emory Kristof, Stephen Low, and Michael V. DeGruy for inspiration and assistance with the procurement of video images using a variety of camera systems; and Matt Tieger for assistance in the generation of video prints. Supported by National Science Foundation Grants OCE-95-29819 and OCE-96-33131.
See also Deep-Sea Ridges, Microbiology. Hydrothermal Vent Ecology Hydrothermal Vent Fauna, Physiology of
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Further Reading Childress JJ and Fisher CR (1992) The biology of hydrothermal vent animals: physiology, biochemistry, and autotrophic symbioses. Oceanography and Marine Biology Annual Review 30: 337--441. Gage JD and Tyler PA (1991) Deep-Sea Biology: A Natural History of Organisms at the Deep-Sea Floor. Cambridge: Cambridge University Press. Lutz RA (2000) Deep sea vents. National Geographic 198(4): 116--127. Shank TM, Fornari DJ, and Von Damm KL (1998) Temporal and spatial patterns of biological community development at nascent deep-sea hydrothermal vents (91N, East Pacific Rise). Deep-Sea Research. 45: 465– 515. Jones ML (ed.) (1985) The Hydrothermal Vents of the Eastern Pacific: An Overview. Bulletin of the Biological Society of Washington, vol. 6. Washington, DC: Biological Society of Washington. Rona PA, Bostrom K, Laubier L, and Smith KL Jr (eds.) (1983) Hydrothermal Processes at Seafloor Spreading Centers. New York: Plenum Press. Tunnicliffe V (1991) The biology of hydrothermal vents: ecology and evolution. Oceanography and Marine Biology Annual Review 29: 319--417. Van Dover CL (2000) The Ecology of Deep-Sea Hydrothermal Vents. Princeton, NJ: Princeton University Press, Princeton.
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HYDROTHERMAL VENT DEPOSITS R. M. Haymon, University of California, CA, USA Copyright & 2001 Elsevier Ltd. All rights reserved. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1228 –1234, & 2001, Elsevier Ltd.
Where Deposits Form: Geologic Controls
Introduction In April 1979, submersible divers exploring the midocean ridge crest at latitude 211N on the East Pacific Rise discovered superheated (3807301C) fluids, blackened by tiny metal-sulfide mineral crystals, spewing from the seafloor through tall mineral conduits (see Hydrothermal Vent Biota, Hydrothermal Vent Fluids, Chemistry of). The crystalline conduits at these ‘black smoker’ hydrothermal vents were made of minerals rich in copper, iron, zinc, and other metals. Since 1979, hundreds of similar hydrothermal deposits have been located along the midocean ridge. It is now clear that deposition of hydrothermal mineral deposits is a common process, and is integrally linked to cracking, magmatism, and cooling of new seafloor as it accretes and spreads away from the ridge (see Propagating Rifts and Microplates, Mid-Ocean Ridge Geochemistry and Petrology, Seamounts and Off-Ridge Volcanism, Mid-Ocean Ridge Seismic Structure). For thousands of years before mid-ocean ridge hot springs were discovered in the oceans, people mined copper from mineral deposits that were originally formed on oceanic spreading ridges. These fossil deposits are embedded in old fragments of seafloor called ‘ophiolites’ that have been uplifted and emplaced onto land by fault movements. The copperrich mineral deposits in the Troodos ophiolite of Cyprus are well-known examples of fossil oceanridge deposits that have been mined for at least 2500 years; in fact, the word ‘copper’ is derived from the Latin word ‘cyprium’ which means ‘from Cyprus.’ The mineral deposits accumulating today at hot springs along the mid-ocean ridge are habitats for a variety of remarkable organisms ranging in size from tiny microbes to large worms (see Hydrothermal Vent Biota, Deep-Sea Ridges, Microbiology, Hydrothermal Vent Fauna, Physiology of). The properties of the mineral deposits are inextricably linked to the organisms that inhabit them. The mineral deposits contain important clues about the physical–chemical environments in which some of these organisms live, and also preserve fossils of some organisms, creating a geologic record of their existence.
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Hydrothermal vent deposits are thus a renewable source of metals and a record of the physical, chemical, biological, and geological processes at modern and ancient submarine vents.
Less than 2% of the total area of the mid-ocean ridge crest has been studied at a resolution sufficient to reveal the spatial distribution of hydrothermal vents, mineral deposits, and other significant small-scale geologic features. Nevertheless, because study areas have been carefully selected and strategically surveyed, much has been learned about where vents and deposits form, and about the geologic controls on their distribution. The basic requirements for hydrothermal systems include heat to drive fluid circulation, and high-permeability pathways to facilitate fluid flow through crustal rocks. On midocean ridges, vents and deposits are forming at sites where ascending magma intrusions introduce heat into the permeable shallow crust, and at sites where deep cracks provide permeability and fluid access to heat sources at depth.
Fast-spreading Ridges Near- and on-bottom studies along the fast-spreading East Pacific Rise suggest that most hydrothermal mineral deposits form along the summit of the ridge crest within a narrow ‘axial zone’ less than 500 m wide. Only a few active sites of mineral deposition have been located outside this zone; however, more exploration of the vast area outside the axial zone is needed to establish unequivocally whether or not mineral deposition is uncommon in this region. The overall spatial distribution of hydrothermal vents and mineral deposits along fast-spreading ridges traces the segmented configuration of cracks and magma sources along the ridge crest (see Propagating Rifts and Microplates, Mid-Ocean Ridge Geochemistry and Petrology, Seamounts and Off-Ridge Volcanism, Mid-Ocean Ridge Seismic Structure). Within the axial zone, mineral deposition is concentrated along the floors and walls of axial troughs created by volcanic collapse and/or faulting along the summit of the ridge crest. The majority of the deposits are located along fissures that have opened above magmatic dike intrusions, and along collapsed
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lava ponds formed above these fissures by pooling and drainage of erupted lava. Where fault-bounded troughs have formed along the summit of the ridge crest, mineral deposition is focused along the bounding faults and also along fissures and collapsed lava ponds in the trough floor. Hydrothermal vents appear to be most abundant along magmatically inflated segments of fast-spreading ridges; however, the mineral deposits precipitated on the seafloor on magmatically active segments are often buried beneath frequent eruptions of new lava flows. The greatest number of deposits, therefore, are observed on inflated ridge segments that are surfaced by somewhat older flows, i.e., along segments where: (1) much heat is available to power hydrothermal vents; and (2) mineral deposits have had time to develop but have not yet been buried by renewed eruptions.
Intermediate- and Slow-spreading Ridges Most hydrothermal deposits that have been found on intermediate- and slow-spreading ridge crests are focused along faults, fissures, and volcanic structures within large rift valleys that are several kilometers wide. The fault scarps along the margins of rift valleys are common sites for hydrothermal venting and mineral deposition. Fault intersections are thought to be particularly favorable sites for hydrothermal mineral deposition because they are zones of high permeability that can focus fluid flow. Mineral deposition on rift valley floors is observed along fissures above dike intrusions, along eruptive fissures and
Dead edifice
Flanges
volcanic collapse troughs, and on top of volcanic mounds, cones and other constructions. In general at slower-spreading ridges, faults appear to play a greater role in controlling the distribution of hydrothermal vents and mineral deposits than they do at fast-spreading ridges, where magmatic fissures are clearly a dominant geologic control on where vents and deposits are forming.
Structures, Morphologies, and Sizes of Deposits A typical hydrothermal mineral deposit on an unsedimented mid-ocean ridge accumulates directly on top of the volcanic flows covering the ridge crest. On sedimented ridges, minerals are deposited within and on top of the sediments. Beneath seafloor mineral deposits are networks of feeder cracks through which fluids travel to the seafloor. Precipitation of hydrothermal minerals in these cracks and in the surrounding rocks or sediments creates a subseafloor zone of mineralization called a ‘stockwork’. In hydrothermal systems where fluid flow is weak, unfocused, or where the fluids mix extensively with sea water beneath the seafloor, most of the minerals will precipitate in the stockwork rather than on the seafloor. Hydrothermal deposits on mid-ocean ridges are composed of: (1) vertical structures, including individual conduits known as ‘chimneys’ (Figure 1) and larger structures of coalesced conduits that are often called ‘edifices’; (2) horizontal ‘flange’ structures that extend outward from chimneys and edifices
Black smoke plume
Fossil worm tubes
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Black smoker chimneys
T decreasing, mixing increasing
White smoke plume White smoker edifices Live alvinelline worm Basal mound
Basalt Figure 1 Composite sketch of the mineral structures and zones in hydrothermal mineral deposits on unsedimented ridge crests (modified after Haymon, 1989). Although mound interiors are seldom observed on the seafloor, the simplified sketch of mineral zoning within the mound is predicted by analogy with chimneys and massive sulfide deposits exposed in ophiolites. An outer peripheral zone (unshaded) of anhydrite þ amorphous silica þ Zn-rich sulfide, dominantly ZnS þ FeS2, is replaced in the interior by an inner zone (hatched) of Cu-rich sulfide (CuFeS2 þ FeS2) þ minor anhydrite and amorphous silica. The inner zone may be replaced by a basal zone (cross-pattern) of Cu-rich sulfide (CuFeS2 þ FeS2) þ quartz. Zones migrate as thermochemical conditions within the mound evolve. Although not shown here, it is expected that zoning around individual fractures cutting through the mound will be superimposed on the simplified zone structure in this sketch.
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at a particular site, which depends on the nature of the heat source and plumbing system, and the rate of seafloor spreading; frequency with which deposits are buried beneath lava flows; and the compositions of the vent fluids and minerals. The large deposits found on slower-spreading ridge crests are located on faults that have moved slowly away from the ridge axis and have experienced repeated episodes of venting and accumulated mineral deposition over thousands of years, without being buried by lava flows. The tall Endeavor Segment edifices are formed because ammonia-enriched fluid compositions favor precipitation of silica in the edifice walls. The silica is strong enough to stabilize these structures so that they do not collapse as they grow taller.
How Do Chimneys Grow? A relatively simple two-stage inorganic growth model has been advanced to explain the basic characteristics of black smoker chimneys (Figure 2). In this model, a chimney wall composed largely of anhydrite (calcium sulfate) precipitates initially from sea water that is heated around discharging jets of hydrothermal fluid. The anhydrite-rich chimney wall precipitated during stage I contains only a small Cross-section chimney wall +
+
+
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Hydrothermal fluid
+ +
+ +
+ +
+ +
+ +
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Pillow basalt (A)
Arrows indicate directions of growth
+ + +
Anhydrite ± (caminite) + FeS2 and Zn(Fe)S Fine-grained pyrrhotite + pyrite + sphalerite intergrowths in an anhydrite matrix +
+
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Hydrothermal fluid
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(Figure 1); (3) mounds of accumulated mineral precipitates (Figure 1); and (4) horizontal layers of hydrothermal sediments, debris, and encrustations. Chimneys are initially built directly on top of the seabed around focused jets of high-temperature effluents. Chimneys and edifices are physically unstable and often break or collapse into pieces that accumulate into piles of debris. The debris piles are cemented into consolidated mounds by precipitation of minerals from solutions percolating through the piles. New chimneys are constructed on top of the mounds as the mounds grow in size. Hydrothermal plume particles and particulate debris from chimneys settle around the periphery of the mounds to form layers of hydrothermal sediment. Diffuse seepage of fluids also precipitates mineral encrustations on mound surfaces, on volcanic flows and sediments, and on biological substrates, such as microbial mats or the shells and tubes of sessile macrofauna. The morphologies of chimneys are highly variable and evolve as the chimneys grow, becoming more complex with time. Black smoker chimneys are often colonized by organisms and evolve into ‘white smokers’ that emit diffuse, diluted vent fluids through a porous carapace of worm tubes (Figure 1). Fluid compositions and temperatures, flow dynamics, and biota are all factors that influence the development of chimney morphology. The complexity of the interactions between these factors, and the high degree of spatial–temporal heterogeneity in the physical, chemical, biological and geological conditions influencing chimney growth, account for the diverse morphologies exhibited by chimneys, and present a challenge to researchers attempting to unravel the processes producing these morphologies. The sizes of hydrothermal mineral deposits on ridges also vary widely. It has been suggested that the largest deposits are accumulating on sedimented ridges, where almost all of the metals in the fluids are deposited within the sediments rather than being dispersed into the oceans by hydrothermal plumes. On unsedimented ridges, the structures deposited on the seabed at fast spreading rates are usually relatively small in dimension (mounds are typically less than a few meters in thickness and less than tens of meters in length, and vertical structures are o15 m high). On intermediate- and slow-spreading ridges, mounds are sometimes much larger (up to tens of meters in thickness, and up to 300 m in length). On the Endeavour Segment of the Juan de Fuca Ridge, vertical structures reach heights of 45 m. The size of a deposit depends on many factors, including: magnitude of the heat source, which influences the duration of venting and mineral deposition; tendency of venting and mineral deposition to recur episodically
Chalcopyrite, or cubanite + (pyrrhotite) Cu_Fe sulfides in an anhydrite matrix As in stage I:higher sulfide:sulfate ratio
Figure 2 Two-stage model of black smoker chimney growth. (A) Stage I, sulfate-dominated stage; (B) stage II, sulfide replacement stage. During stage II, several different sulfide mineral zonation sequences develop, depending on permeability and thickness of chimney walls, hydrodynamic variables, and hydrothermal fluid composition.
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component of metal sulfide mineral particles that crystallize because of rapid chilling of the hydrothermal fluids. In stage II, the anhydrite-rich wall continues to grow upward and to thicken radially,
Table 1 Minerals occurring in ocean ridge hydrothermal mineral deposits
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protecting the fluid flowing through the chimney from very rapid chilling and dilution by sea water. This allows metal sulfide minerals to precipitate into the central conduit of the chimney from the hydrothermal fluid. The hydrothermal fluid percolates outward through the chimney wall, gradually replacing anhydrite and filling voids with metal sulfide minerals. During stage II, the chimney increases in height, girth and wall thickness, and both the calcium sulfate/metal sulfide ratio and permeability of the walls decrease. Equilibration of minerals with pore fluid in the walls occurs continuously along steep, time–variant temperature and chemical gradients between fluids in the central conduit and sea water surrounding the chimney. This equilibration produces sequences of concentric mineral zones across chimney walls that evolve with changes in thermal and chemical gradients and wall permeability. The model of chimney growth described above is accurate but incomplete, as it does not include the effects on chimney development of fluid phase separation, biological activity, or variations in fluid composition. Augmented models that address these
Table 2 Ranges of elemental compositions in bulk midocean ridge hydrothermal mineral deposits
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Morphological and Mineralogical Evolution of Chimneys on the East Pacific Rise at 9˚-10˚N Stage 2
1991 "Proto-chimney" Anhydrite-dominated T = 389˚-403˚C
Stage 2
1992-1995 CuFe-sulfide-dominated T = 340˚-392˚C
1992-1995 Zn-sulfide-dominated T = 264˚-340˚C
KEY an _ anhydrite cp _ chalcopyrite po _ pyrrhotite py _ pyrite
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HYDROTHERMAL VENT DEPOSITS
complexities are needed to fully characterize the processes governing chimney growth.
Elemental and Mineral Compositions of Deposits Ridge crest hydrothermal deposits are composed predominantly of iron-, copper- and zinc sulfide minerals, calcium- and barium-sulfate minerals, iron oxide and iron oxyhydroxide minerals, and silicate minerals (Table 1). These minerals precipitate from diverse processes, including: heating of sea water; cooling of hydrothermal fluid; mixing between sea water and hydrothermal fluid; reaction of hydrothermal minerals with fluid, sea water, or fluid–sea water mixtures; reaction between hydrothermal fluid and seafloor rocks and sediments; and reactions that are mediated or catalyzed biologically. This diversity in the processes and environments of mineral precipitation results in the deposition of many different minerals and elements (Tables 1 and 2). High concentrations of strategic and precious metals are found in some deposits (Table 2). The deposits are potentially valuable, if economic and environmentally safe methods of mining them can be developed. Chimneys can be classified broadly by composition into four groups: sulfate-rich, copper-rich, zinc-rich and silica-rich structures. Copper-rich chimney compositions are indicative of formation at temperatures above 3001C. Sulfate-rich compositions are characteristic of active and immature chimneys. Many chimneys are mineralogically zoned, with hot interior regions enriched in copper, and cooler exterior zones enriched in iron, zinc, and sulfate (Figures 1 and 2). Mounds exhibit a similar gross mineral zoning, and those which are exposed by erosion in ophiolites often have silicified (quartzrich) interiors (Figure 2). Seafloor weathering of deposits after active venting ceases results in dissolution of anhydrite, and oxidation and dissolution of metal-sulfide minerals. Small deposits that are not sealed by silicification or buried by lava flows will
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not be well preserved in the geologic record (Figure 3).
Chimneys as Habitats Chimney and mound surfaces are substrates populated by microbial colonies and sessile organisms such as vestimentiferan and polychaete worms, limpets, mussels, and clams. It is likely that pore spaces in exterior regions of chimney walls are also inhabited by microbes. All of these organisms that are dependent on chemosynthesis benefit from the seepage of hydrothermal fluid through active mineral structures, and from the thermal and chemical gradients across mineral structures. The structures provide an interface between sea water and hydrothermal fluid that maintains tolerable temperatures for biota, and allows organisms simultaneous access to the chemical constituents in both sea water and hydrothermal fluid. However, organisms attached to active mineral structures must cope with changes in fluid flow across chimney walls (which sometimes occur rapidly), and with ongoing engulfment by mineral precipitation. Some organisms actively participate in the precipitation of minerals; for example, sulfide-oxidizing microbes mediate the crystallization of native sulfur crystals, and microbes are also thought to participate in the precipitation of marcasite and iron oxide minerals. Additionally, the surfaces of organisms provide favorable sites for nucleation and growth of amorphous silica, metal sulfide and metal oxide crystals, and this facilitates mineral precipitation and fossilization of vent fauna (Figure 3).
Fossil Record of Hydrothermal Vent Organisms Fossil molds and casts of worm tubes, mollusc shells, and microbial filaments have been identified in both modern ridge hydrothermal deposits and in Cretaceous, Jurassic, Devonian, and Silurian deposits. This fossil record establishes the antiquity of vent
Figure 3 (Left)On left: a time series of seafloor photographs showing the morphological development of a chimney that grew on top of lava flows erupted in 1991 on the crest of the East Pacific Rise near 9150.30 N (Haymon et al., 1993). Within a few days-to-weeks after the eruption, anhydrite-rich ‘Stage 1 Protochimneys’ a few cm high had formed where hot fluids emerged from volcanic outcrops covered with white microbial mats (top left). Eleven months later, the chimney consisted of cylindrical ‘Stage 2’ anhydrite-sulfide mineral spires approximately one meter in height, and as-yet unpopulated by macrofauna (middle left). Three and a half years after the eruption, the cylindrical conduits had coalesced into a 7 m-high chimneys structure that was covered with inhabited Alvinelline worms tubes (bottom left). On right: photomicrographs of chimney samples from the eruption area that show how the chimneys evolved from Stage 1 (anhydrite-dominated; top right) to Stage 2 (metal-sulfide dominated) mineral compositions (see text). As the fluids passing through the chimneys cooled below B3301C during Stage 2, the CuFe-sulfide minerals in the chimney walls (middle right) were replaced by Zn- and Fe-sulfide minerals (bottom right).
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communities and the long evolutionary history of specific faunal groups. The singular Jurassic fossil assemblage preserved in a small ophiolite-hosted deposit in central California is particularly interesting because it contains fossils of vestimentiferan worms, gastropods and brachiopods, but no clam or mussel fossils. In contrast, modern and Paleozoic faunal assemblages described thus far include clams, mussels and gastropods, but no brachiopods. Does this mean that brachiopods have competed with molluscs for ecological niches at vents, and have moved in and out of the hydrothermal vent environment over time? Fossilization of organisms is a selective process that does not preserve all the fauna that are present at vents. Identification of fossils at the species level is often difficult, especially where microbes are concerned. Notwithstanding, it is important to search for more examples of ancient fossil assemblages and to trace the fossil record of life at hydrothermal vents back as far as possible to shed light on how vent communities have evolved, and whether life on earth might have originated at submarine hydrothermal vents.
Summary Formation of hydrothermal deposits is an integral aspect of seafloor accretion at mid-ocean ridges. These deposits are valuable for their metals, for the role that they play in fostering hydrothermal vent ecosystems, for the clues that they hold to understanding spatial–temporal variability in hydrothermal vent systems, and as geologic records of how life at hydrothermal vents has evolved. From these deposits we may gain insights about biogeochemical processes at high temperatures and pressures that can be applied to understanding life in inaccessible realms within the earth’s crust or on other planetary bodies. We are only beginning to unravel the complexities of ridge hydrothermal vent deposits. Much exploration and interdisciplinary study remains to be done to obtain the valuable information that they contain.
See also Deep-Sea Ridges, Microbiology. Hydrothermal Vent Fluids, Chemistry of. Hydrothermal Vent Biota. Hydrothermal Vent Fauna, Physiology of. Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Seismic Structure. Propagating Rifts and Microplates. Seamounts and Off-Ridge Volcanism.
Further Reading Dilek Y, Moores E, Elthon D, and Nicolas A (eds.) (2000) Ophiolites and Oceanic Crust: New Insights from Field Studies and the Ocean Drilling Program. Geological Society of America Memoir. Boulder: Geological Society of America. Haymon RM (1989) Hydrothermal processes and products on the Galapagos Rift and East Pacific Rise, 1989. In: Winterer EL, Hussong DM, and Decker RW (eds.) The Geology of North America: The Eastern Pacific Ocean and Hawaii, vol. N, pp. 125--144. Boulder: Geological Society of America. Haymon RM (1996) The response of ridge crest hydrothermal systems to segmented, episodic magma supply. In: MacLeod CJ, Tyler P, and Walker CL (eds.) Tectonic, Magmatic, Hydrothermal, and Biological Segmentation of Mid-Ocean Ridges, vol. Special Publication 118, pp. 157--168. London: Geological Society. Humphris SE, Zierenberg RA, Mullineaux LS, and Thomson RE (eds.) (1995) Seafloor Hydrothermal Systems: Physical, Chemical, Biological, and Geological Interactions, Geophysical Monograph, vol. 91. Washington, DC: American Geophysical Union. Little CTS, Herrington RJ, Haymon RM, and Danelian T (1999) Early Jurassic hydrothermal vent community from the Franciscan Complex, San Rafael Mountains, California. Geology 27: 167--170. Tivey MK, Stakes DS, Cook TL, Hannington MD, and Petersen S (1999) A model for growth of steep-sided vent structures on the Endeavour Segment of the Juan de Fuca Ridge: results of a petrological and geochemical study. Journal of Geophysical Research 104: 22859--22883.
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HYDROTHERMAL VENT ECOLOGY C. L. Van Dover, The College of William and Mary, Williamsburg, VA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1234–1241, & 2001, Elsevier Ltd.
Introduction Most of the ocean floor is covered with a thick layer of sediment and is populated by sparse and minute, mud-dwelling and mud-consuming invertebrates. In striking contrast, the volcanic basalt pavement of mid-ocean ridges hosts hydrothermal vents and their attendant lush communities of large invertebrates that ultimately rely on inorganic chemicals for their nutrition. Vents themselves are sustained by tectonic forces that fracture the basalt and allow sea water to penetrate deep within the ocean crust, and by volcanism, which generates the hot rock at depth that strips sea water of oxygen and magnesium. The hot rock gives up to the nascent vent fluid a variety of metals, especially copper, iron, and zinc, as well as reduced compounds such as hydrogen sulfide and methane. The vent fluid, thermally buoyant, rises to exit as hot springs on the seafloor. Discovered first by geologists in 1977 along a stretch of mountain range known as the Galapagos Rift, near the equator in the eastern Pacific Ocean, hydrothermal vents are now known to occur along every major ridge system on the planet. Several of these ridge systems – in the Arctic and Antarctic, in the Indian Ocean, in the southern Atlantic – are only just beginning to be explored. Regional species composition of the vent fauna differs between ocean basins and, sometimes, even within a basin. Most of the species that occur at vents have never been found in the adjacent, non-vent deep sea and are considered to be endemic, adapted to the chemical milieu of the vent environment. Reduced compounds carried in hydrothermal fluids, together with oxygen from the surrounding sea water, fuel the microbial fixation of inorganic carbon into organic carbon that forms the chemosynthetic base of the vent food web. Hydrothermal vents on midocean ridges are thus globally distributed, insular ecosystems that support endemic faunas through chemosynthetic processes rather than through photosynthesis. They are effectively decoupled by depth (typically 41000 m) from climatic variations and anthropogenic activities, but are tightly coupled to geophysical processes
of tectonism and volcanism. Vents thus offer unique opportunities for biologists to study adaptations that allow life to persist in this extreme environment and to explore planetary controls on biodiversity and biogeography along submarine, hydrothermal ‘archipelagoes’, where propagules are water-borne and subject to dispersal in an open system. Further, because vents are thought to have been a primordial component of the oceans and because there is increasing speculation that early life on this planet may have thrived in hot environments and on chemicals rather than on an organic soup for nourishment, extant hydrothermal systems are thought by many to be analogues for sites where early life may have evolved on this and other planets or planetary bodies in our solar system.
Microorganisms and the Chemosynthetic Basis for Life at Vents The terms ‘chemosynthesis’ and ‘photosynthesis’ are imprecise. While a voluminous nomenclature is available to differentiate among variations in these processes, for simplicity, chemosynthesis and photosynthesis are used here. In photosynthesis, sunlight captured by proteins provides energy for the conversion of inorganic carbon (carbon dioxide, CO2) and water (H2O) into organic carbon (carbohydrates, [CH2O] and oxygen (O2) (eqn [1]). light
CO2 þ H2 O - ½CH2 O þ O2
½1
Photosynthesis by plants is the basis for consumer and degradative food webs both on land and, as a rain of organic detritus derived from surface phytoplankton productivity, on the seabed. In the deep sea, detrital inputs of organic carbon are exceedingly small, accounting for the paucity of consumer biomass in abyssal muds. At hydrothermal vents, the supply of surface-derived organic material is overwhelmed by the supply of new organic carbon generated through chemical oxidation of hydrogen sulfide (H2S) (eqn [2]). CO2 þ H2 O þ H2 S þ O2 -½CH2 O þ H2 SO4
½2
Metabolic fixation pathways for carbon can be identical in photosynthetic plants and chemosynthetic
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microorganisms, namely the Calvin–Benson cycle, but the energy-yielding processes that fuel the Calvin– Benson cycle (photon capture versus chemical oxidation) are distinctive. High biomass at hydrothermal vents is in part a consequence of the aerobic nature of the process described in eqn [2]. Oxygen is used to oxidize the hydrogen sulfide, generating a large energy yield that in turn can fuel the production of large
amounts of organic carbon (Figure 1). Nonaerobic chemical reactions, such as oxidation of vent-supplied hydrogen (H2) by carbon dioxide (CO2), can also support chemosynthesis at vents, but energy yields under such anaerobic conditions are much lower than from aerobic oxidation. Microorganisms using these anaerobic reactions cannot by themselves support complex food webs and large invertebrates.
Symbiosis and the Host–Symbiont Relationship One of the hallmarks of many hydrothermal vent communities is the dominance of the biomass by invertebrate species that host chemosynthetic microorganisms within their tissues. Giant, redplumed, vestimentiferan tubeworms (Riftia pachyptila; Figure 2) so far provide the ultimate in host accommodation of endosymbiotic bacteria. These worms live in white, chitinous tubes, with their plumes extended into the zone of turbulent mixing of warm (B201C), sulfide-rich, hydrothermal fluid and cold (21C), oxygenated sea water. When discovered
CO2+H2O solar energy [CH2O] + O2 Photosynthesis
Chemical energy CO2 + H2O + H2S + O2 [CH2O] + H2SO4
Chemosynthesis
Figure 1 Photosynthetic and chemosynthetic processes in the ocean. Sunlight fuels the generation of organic material (CH2O) from inorganic carbon dioxide (CO2) and water (H2O) by phytoplankton in surface, illuminated waters. At depths where hydrothermal vents exist (typically>2000 m), no sunlight penetrates. In place of sunlight, the chemical oxidation of sulfide (H2S) by oxygen (O2) fuels the conversion of carbon dioxide to organic carbon by chemosynthetic bacteria.
Figure 2 The giant tubeworm Riftia pachyptila. The red plume of the tubeworm acts as a gill for uptake of dissolved gases. The trunk of the worm is found inside the white, chitinous tube. (Photograph by Dudley Foster, Woods Hole Oceanographic Institution.)
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in 1977, vestimentiferan tubeworms were remarkable for their size (up to several meters in length) and the complete absence of a digestive system in adults. In fact, the digestive system has been replaced by the trophosome, which is a specialized, paired organ derived from the larval gut. The trophosome is richly infiltrated with blood capillaries and each of its lobes lies within a blood-filled body cavity. The blood itself is rich in hemoglobin. Host bacteriocyte cells in the trophosome house chemosynthetic bacteria that use hydrogen sulfide and oxygen to fuel the production of organic carbon, as described above. The metabolic requirements of the tubeworm endosymbiotic bacteria place some remarkable burdens on the host. First, there is a novel requirement for delivery of sulfide to the bacteria, which reside at a location remote from the site of gas exchange (the plume). Sulfide is normally a potent toxin to animals, poisoning the cellular enzyme system that generates ATP (adenosine triphosphate), the currency of metabolism. Sulfide also competes with oxygen for binding sites on hemoglobin. Tubeworm hemoglobin has separate binding sites for oxygen and sulfide, so that both can be transported throughout the worm in the circulatory system without competition. When bound to hemoglobin, the sulfide is not reactive and so enzyme systems remain unchallenged. Once delivered to the bacteria in the trophosome, the sulfide is quickly oxidized and loses its toxic potential. Novel requirements for carbon dioxide are also found in tubeworms. The usual flow of CO2 is out of an animal, as the end-product of metabolism, but the resident bacteria of the trophosome require a net uptake of CO2. Maintenance of high concentrations of inorganic carbon in the blood of the tubeworm is facilitated by the high partial pressure of CO2 in the water surrounding the site of uptake (the plume) and by the alkaline internal pH of the blood (7.3–7.4), which favors the bicarbonate form (HCO 3 ) of carbon dioxide and thus maintains a steep concentration gradient for diffusion of CO2 from the environment into the blood. As described above, the anatomy of the tubeworm is well adapted for life in sulfide- and CO2-rich vent fluids and for supporting its endosymbiotic, chemosynthetic bacteria. The bacteria provide nearly all of the nutrition for the host, with the exception, perhaps, of small amounts of dissolved organic materials taken up across the tissues of the plume. In turn, the bacteria are provided with a chemically rich and stable environment for growth. Other large invertebrates at vents also derive much of their nutrition from endosymbiotic, chemosynthetic bacteria, including 20 to 30-cm long to vesicomyid clams and bathymodiolid mussels. While
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Figure 3 Swarming shrimp (Rimicaris exoculata) at a hydrothermal vent on the Mid-Atlantic Ridge. (Photograph by C.L. Van Dover.)
vent mussels have a fairly normal digestive system and are capable of filter-feeding just as shallow-water mussels do, vent clams lack a functional digestive system. Both types of bivalves have enlarged gills and it is within these gills that the endosymbiotic bacteria are found. Clams are thought to take up hydrogen sulfide via their highly vascularized foot, with which they probe cracks in the basalt where vent fluids emanate. Not all chemosynthetic bacteria that nourish vent invertebrates are endosymbiotic. Shrimp that dominate vents in the Atlantic (Figure 3) host chemosynthetic bacteria on their carapace (i.e., the bacteria are episymbiotic) and seem to depend on these bacteria for a significant portion of their diet. Other chemosynthetic bacteria are free-living, suspended in the water column, providing nourishment to suspension-feeding invertebrates such as barnacles, or grow as mats or films on surfaces, where grazers such as limpets and polychaetes forage. Heterotrophic bacteria (using organic rather than inorganic compounds) may also be important for consumers within the vent invertebrate food web, but this has yet to be examined carefully.
Thermal Adaptations While hydrothermal vent communities live at temperatures slightly elevated above ambient sea water temperature, the existence of highly productive communities at vents is a consequence of fluid chemistry rather than thermal input. Nevertheless, there are some invertebrates, notably the large, thumb-sized polychaetes in the family Alvinellidae (Figure 4), that are especially tolerant of high temperatures and that compete with desert ants for the
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Guinness Book of World Records in the category of multicellular animal living at a thermal extreme. Alvinellids live on the sides of black smoker sulfide chimneys, and are reported to survive brief exposures to temperatures as high as 1051C. They routinely experience a thermal gradient of 50–601C over the length of their bodies. Temperature tolerance in these worms is not completely understood, but thermal stability of enzymes has been shown to increase in alvinellid species that occupy increasingly warmer habitats. Membranes may also be adapted for thermal stability through an increase in the degree of double bonding in fatty acids in species living in warmer portions of the vent habitat. Shrimp (Rimicaris exoculata) that swarm on black smoker chimneys in close proximity to extremely high temperature (3501C) fluids are not bathed in excessively hot water, but they do have paired photoreceptive organs that may function to detect the glow emitted by the hot water. The organs are derived from ordinary shrimp eyes, but the photoreceptive surfaces are hypertrophied and rich in visual pigment (rhodopsin). The eyestalks are lost and the derived eyes extend back along the dorsal surface of the shrimp, beneath the transparent carapace. These ‘eyes’ have also lost their optic (lens) systems, so they cannot form an image, but they are
Figure 4 The alvinellid polychaete Alvinella caudata, beside its fragile tube. (Photograph by J. Porteous, Woods Hole Oceanographic Institution.)
optimized to detect gradients of dim light. While light from black smokers has now been well documented, the behavioral response of the shrimp to this light has not been studied.
Community Dynamics There could hardly be a greater dynamic contrast than that found between the cold, food-limited, relentlessly stable and vast deep sea environment, and the thermally complex, trophically rich, ephemeral, and insular deep-sea hydrothermal vent fields. Since the discovery of vents, ecologists have attempted to predict the cycle of community development over the life span of a vent field by interpolation and extrapolation from snapshot observations. The 1991 seafloor volcanic eruption at the Venture Hydrothermal Field on the East Pacific Rise was witnessed within days to weeks of the event, providing biologists with the first submarine equivalent of Krakatau. At the time of the Venture eruption, existing vent communities were obliterated by fresh lava flows, pervasive warm-water venting was observed along B1.5 km of ridge axis, and a dense ‘bloom’ of flocculent material poured from cracks between lobes of lava, obscuring visual navigation (Figure 5). The flocs were determined to be the mineral by-product of microbial production. Within one year, venting became focused at numerous sites along the ridge axis and the first colonists had arrived, including a small tubeworm species (Tevnia jerichonana) and dense aggregations of several species of limpets. Mobile vent organisms, including vent crabs and squat lobsters, zoarcid fish, and swarming amphipods were also well represented. After 2.5 years, some vents had shut off but, where venting persisted, mature colonies of the giant tubeworm (Riftia
Figure 5 Flocculent material suspended in the water column 1 m above new ocean crust. The floc is a mineral by-product of microbial production and emanated from cracks in the seafloor. (Photograph by Alvin, Woods Hole Oceanographic Institution.)
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pachyptila) were established, along with a variety of smaller invertebrates that live among the tubeworm tubes, including shrimp, limpets, amphipods, and polychaetes. Growth rates of R. pachyptila were measured to be among the most rapid of any aquatic invertebrate. Within 5 years, 75% of the regional species pool could be found at the new vents and mussel beds were well-established and beginning to overwhelm the R. pachyptila thickets. In this example, vesicoymid clams were the last of the big megafaunal species to arrive at the site, despite the presence of adult populations within several kilometers of the area overrun by fresh lava. Mussels appear to have a competitive edge among the larger taxa at vents, in part because they are mobile and can relocate as necessary to cope with changing flow patterns of vent fluids, while tubeworms are stationary and have few options for tracking vent flow. In addition, because mussels can filter-feed as well as derive nutrition from their endosymbiotic bacteria, they are among the last species to disappear as a vent shuts down. Ultimately, it is the mobile scavengers – the crabs and fish and octopus – that witness the final demise of a vent field. The eruption at the Venture Hydrothermal Field was not entirely unexpected – geologists were studying this region of the ridge axis because it was so inflated and appeared to be ripe for an eruptive event. Localization of seafloor eruptions on the Juan de Fuca Ridge in the northeast Pacific (off the coast of Vancouver Island) now takes place in real-time, facilitated by a legacy of the cold war era, namely through acoustic signals received by underwater sound-surveillance systems originally designed to track enemy submarines. As eruptions take place, T-waves (low-frequency, tertiary waves characteristic of lava in motion) are transmitted into the water column. Navy hydrophone networks allow the T-waves to be placed in a geographical context and traced from start to finish. One migration of lava in June 1993 was traced for about 40 km below the ocean crust over a two-day period before it finally erupted onto the seafloor. As at the Venture Field, fresh lava on the seafloor was observed, along with venting of flocculent material derived from microbial production. Sites of persistent venting were colonized by populations of vent invertebrates within one year, some of which were reproductively mature. The two examples of community dynamics cited here suggest that the species composition of all vent communities is constantly changing. But some vent sites are long-lived, and repeat visits to these longlived sites over a 15-year period document essentially no change in the nature of the fauna. The best
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example of this to date is the TAG site on the MidAtlantic Ridge. TAG is a large sulfide mound (100 m diameter) that has occupied the junction of crosscutting faults and fissures for more than 100 000 years. Swarming shrimp (Rimicaris exoculata) and anemones (Maractis rimicarivora) dominated the site when it was first discovered in 1985 and continued to dominate through the most recent set of observations in 1998, despite massive disruptions of hydrothermal flow caused by drilling in 1994. There have been shifts in the precise location of the primary masses of shrimp as they track local natural and anthropogenic changes in vent flow on the mound, and we assume that there has been replacement of generations by recruitment. But there has been no succession observed, no invasion by other taxa. The absence of change does little to attract ecologists, who thrive on dynamic systems, but the stability of the species composition at TAG and other sites has profound implications regarding the selective pressures encountered by the species that inhabit these sites compared to sites that are constantly threatened by lava overruns or tectonic shifts in plumbing.
Origins of Vent Faunas, Biogeography, and Biodiversity Evolutionary paths that brought invertebrate taxa to vents are varied. Several major taxa, including the galatheid squat lobsters, pycnogonid sea spiders, and echinoderms are likely to be immigrants from the surrounding deep sea. Some species are closely related to (and presumably derived from) shallowwater genera. Many of the most familiar vent taxa – the vesicomyid clams, bathymodiolid mussels, vestimentiferan tubeworms, alvinocarid shrimp – are allied to genera and families found in a variety of deepsea chemosynthetic ecosystems (i.e., seeps and whale-falls as well as vents). The direction of invasion (seep-to-vent or vice versa) can be inferred using molecular techniques. For example, molecular phylogenetics suggests that the bathymodiolid mussel group invaded vents from seeps, with several seep species resulting from reinvasion of the seep habitat by vent ancestors. There are also specialized taxa so far known only from hydrothermal vents. The most conspicuous of these is the alvinellid polychaete family, whose members often occupy the warmest habitable waters of a vent site. Still other taxa appear to be relicts of ancient lineages that have found refuge in the vent environment. These relict taxa include the stalked
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1 cm
Figure 6 The ancient barnacle Neolepas zevinae. (Photograph by C.L. Van Dover.)
Present
barnacle, Neolepas zevinae (Figure 6) and the archaeogastropod limpet, Neomphalus fretterae. Fossil vent communities found on land provide a glimpse of Silurian vent assemblages. Vestimentiferans are reported from some of the oldest vent deposits known, accompanied by brachiopods and monoplacophorans. Brachiopods and monoplacophorans are so far poorly represented in modern vent communities, if at all. When vents were first discovered, some biologists hypothesized that the vent fauna would be globally cosmopolitan. Subsequent explorations have shown this hypothesis to be false, and mechanisms that allow isolation and differentiation of faunas have been postulated. One of the best examples comes from a comparison of faunas from the Juan de Fuca Ridge and East Pacific Rise. At one time (56 Ma), these ridges were one continuous ridge system, but around 37 Ma the North American Plate began to
56 Ma
Kula North America
Pacific Pacific
65 Ma
Farallon
110 Ma
Kula Farallon
Izanagi Farallon Pacific
Pacific
Figure 7 Bisection of a mid-ocean ridge by the North American Plate. At one time (56 Ma), there was a continuous mid-ocean ridge in the eastern Pacific basin but, as the North American Plate overrode the ridge system, it was bisected to form the northeast Pacific ridge system (Juan de Fuca, Explorer, Endeavour, and Gorda Ridges) and the East Pacific Rise (emerging in the Gulf of California and running south toward Antarctica). (Reproduced from Tunnicliffe et al. 1996.)
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override this ridge (Figure 7), splitting the parent faunal assemblage into two daughter assemblages that are distinct yet closely allied at the generic level. Similarities among faunas at the species level may be reduced to nearly zero at vent sites that are in separate ocean basins. This observation suggests that major additions to the global vent faunal inventory await us in the unexplored ocean basins; the potential for discovery of novel adaptations to the vent ecosystem is extremely high. Vent biologists are just beginning to examine global patterns in biodiversity. Preliminary measures show a strong correlation between spreading rate and species richness, raising the hypothesis that patterns of volcanism may be an ultimate control on species diversity in hydrothermal vent ecosystems. Where ridges are fast-spreading, the magma budget is high and the temporal and spatial frequency of vents along the ridge axis is high. In contrast, on slow-spreading ridges, the magma budget is low and vents are far apart. Because distances between vents are short on fast-spreading ridges, species are less likely to go extinct. At slow-spreading ridges, allopatric speciation by isolation may be favored, but species are more likely to go extinct because of the distance between vents. Many of the species that dominate vents on slow-spreading systems have mobile rather than sessile adults, suggesting that distance may act as a selective filter for dispersal capability in these systems.
Vent Systems and the Origin of Life Without doubt, the most provocative consequence of the discovery of seafloor hydrothermal vents is the suggestion that vents may have been the site where life originated on Earth. In contrast to the heterotrophic hypothesis of the origin of life, with its nourishing ‘organic soup’, the vent theory suggests that the earliest life was chemosynthetic, taking biochemical advantage of the large degree of chemical disequilibrium associated with mixing zones of low- and high-temperature portions of hydrothermal systems. A case has been made that the thermophilic nature of the most ancient known lineages of life is indicative of an origin of life at hot springs, but this argument finds limited support, since these most ancient cell types are still very complex and far removed from the progenitors of life. One model for the origin of life at vents suggests the following acellular precursor: a monomolecular, negatively-charged organic layer bonded to positively charged mineral surfaces at the interface of hot water. In this model, pyrite, which forms exergonically from
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iron monosulfide and hydrogen sulfide (both components of vent fluids), serves as the mineral surface. The pyrite-forming reaction yields two free electrons that can be used for building biochemical constituents; simple organic molecules interacting with the pyrite could be reduced to more complex organic molecules. The simple organics also derive from the vent fluids, through abiogenic synthesis. In a secondary stage of development, the precursor evolves to being semicellular, still supported by minerals, but with a lipid membrane and internal broth, with increasing metabolic capabilities. In the final stage of origin, the pyrite support is abandoned and true cellular organisms arise. The elegance of this model is that it uses an energetically realistic inorganic chemical reaction to create a cationic substrate that can bind with organic compounds, all within a single setting. Discovery of vent ecosystems and the appreciation of their chemosynthetic basis has influenced the search for life elsewhere in the solar system. When the Viking Mission to Mars took place, the emphasis was on a search for photoautotrophic processes, but now the search for evidence of past or extant life on other planets highlights environments where chemosynthetic processes may take place, including hydrothermal areas.
Closing Remarks Hydrothermal vent ecology remains a field ripe for discovery of novel faunas and adaptations, as vents in new ocean basins are explored. Because access to deep-sea ecosystems is continually improving, biologists can now undertake quantitative sampling and time-series investigations that are certain to reshape our understanding of the physiological ecology, population biology, community dynamics, and biogeography of vent faunas in the near future.
See also Hydrothermal Vent Biota. Hydrothermal Vent Fluids, Chemistry of. Hydrothermal Vent Deposits. Hydrothermal Vent Fauna, Physiology of. MidOcean Ridge Tectonics, Volcanism, and Geomorphology.
Further Reading Bock GR and Goode JA (eds.) (1996) Evolution of Hydrothermal Ecosystems on Earth (and Mars?). New York: Ciba Foundation.
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Childress JJ and Fisher CR (1992) The biology of hydrothermal vent animals: physiology, biochemistry, and autotrophic symbioses. Oceanography and Marine Biology Annual Review 30: 337--441. Desbruye`res D and Segonzac M (eds.) (1997) Handbook of Deep-Sea Hydrothermal Vent Fauna. Brest: IFREMER. Fisher CR (1990) Chemoautotrophic and methanotrophic symbioses in marine invertebrates. Critical Reviews in Aquatic Science 2: 399--436. Humphris SE, Zierenberg RA, Mullineaux LS, and Thomson RE (eds.) (1995) Seafloor Hydrothermal Systems: Physical, Chemical, Biological and Geological Interactions. Washington, DC: American Geophysical Union. Karl DM (ed.) (1995) The Microbiology of Deep-Sea Hydrothermal Vents. New York: CRC Press. Shank TM, Fornari DJ, Von Damm KL et al. (1998) Temporal and spatial patterns of biological community
development at nascent deep-sea hydrothermal vents (91N, East Pacific Rise). Deep-Sea Research 45: 465--516. Tunnicliffe V (1991) The biology of hydrothermal vents: ecology and evolution. Ocenography and Marine Biology Annual Review 29: 319--407. Tunnicliffe V, Fowler CMR, and McArthur AG (1996) Plate tectonic history and hot vent biogeography. In: MacLeod CJ, Tyler PA, and Walker CL (eds.) Tectonic, Magmatic, Hydrothermal and Biological Segmentation of Mid-Ocean Ridges, Geological Society Special Publication 118, pp. 225–238. Tyler PA and Young CM (1999) Reproduction and dispersal at vents and cold seeps: a review. Journal of the Marine Biology Association of the UK 79: 193--208. Van Dover CL (2000) The Ecology of Deep-Sea Hydrothermal Vents. Princeton: Princeton University Press.
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HYDROTHERMAL VENT FAUNA, PHYSIOLOGY OF A. J. Arp, Romberg Tiburon Center for Environment Studies, Tiburon, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1242–1246, & 2001, Elsevier Ltd.
Hydrothermal Vent Environments are Dynamic, Hot, and Toxic The hydrothermal vent environments, lying at the bottom of the ocean at depths of 2.5 km or more, were discovered in 1977 by a group of geologists exploring spreading centers at midocean ridges on the sea floor. As fissures open up in the earth’s surface, lava is extruded onto the ocean floor and sea water is pulled towards the center of the earth so deeply that it comes into contact with hot, molten magma. The sea water is superheated and then discharged back into the environment through fissures in the ocean floor. As sea water moves from the center of the earth and into the vent habitat on the seafloor, it becomes laden with inorganic chemicals. In particular, hydrogen sulfide, an essential chemical in this unique environment, leaches into the water at depth, and water discharging into the environment can be highly enriched in this toxic but energy-rich chemical. There are two types of hydrothermal vents. In those characterized by diffuse venting, sea water percolates out at a moderate rate and is approximately 10–201C in temperature. As the bottom water temperature of the majority of the earth’s deep ocean is about 21C, these hydrothermal fluids are elevated in temperature, but rapidly mix with the surrounding sea water. One of the first hydrothermal vent environments discovered, Rose Garden at the hydrothermal vent environment near the Galapagos Islands, is an example of a diffuse vent habitat. In the black smoker environment of the hydrothermal vents, things are a lot hotter, such as at those on the Juan de Fuca Ridge off the coast from the state of Washington. This is a very dynamic, high-temperature environment where water issuing forth from large chimneylike structures can be as hot as 4001C in temperature. In spite of the hydrogen sulfide-enriched environment, elevated temperatures, and the dynamic volcanic activity, numerous and varied animals cluster around these sites, taking advantage of the hard substrate provided by the extruded pillow lava. A
typical vent environment begins with a stretch of pillow lava in the foreground, with clams wedged into the numerous fissures. At the heart of the vent environment are diffuse venting hydrothermal fluids or actively spewing white smoker chimneys and black smoker chimneys, teeming with life. Typical inhabitants include dense clusters of tubeworms and many free-ranging animals roaming in and out of the vent environment such as brachyuran crabs, galatheid crabs, numerous amphipods, a few species of fish, and a host of other smaller animals.
Chemosynthesis — The Basis of All Life in the Vent Environment Possibly the most revolutionary outcome of the discovery of hydrothermal vents is the story of how life exists in this challenging habitat and the unique nature of the food chain and the source of basic energy in this remote location. Prior to the discovery of the hydrothermal vents, most biologists believed that all life depended upon the energy of sunlight and that the basis of all food chains was photosynthesis. When the hydrothermal vents were discovered, it was immediately clear that they represent a very enriched biological environment very remote from the surface sunlight. It was difficult to imagine that organic material could drift down in large enough quantities to provide the energy to fuel this environment. Other interesting data materializing rapidly after the discovery of the vents indicated that most of the animals, especially the large invertebrates, have no digestive systems. For example, the large tubeworm, Riftia pachyptila, has no mouth and no intestine; however, it is a large animal approximately 1.5 m in length and up to 2 cm in diameter, and large colonies of these animals flourish in the remote vent habitat. The question of how animal life is supported in the deep sea communities of the hydrothermal vents became, and remains to this day, a focus of intense research. Several lines of evidence lead to the realization that some of the major invertebrates endemic to the vent environment harbor bacteria within their body cavities. Free-living bacteria in this environment and in our own backyard have been known for years to be able to use chemical energy as a basis of their metabolism. In the case of the free-living bacteria, there are many hydrogen sulfide-oxidizing bacteria that can use this chemical as the basis of
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their metabolic pathways and produce organic compounds that form the basis of their nutrition. Hydrothermal vent animals harboring these chemical-utilizing bacteria within their body tissues as symbionts include the large tubeworm Riftia pachyptila, which has dense aggregations of bacteria in a residual gutlike organ, and the clam Calyptogena magnifica, which harbors bacteria in the gills. These symbiotic bacteria are able to utilize the inorganic chemical hydrogen sulfide, so plentiful in this environment, in a manner analogous to what plants do with energy from the sun. This process is therefore termed chemosynthesis rather than photosynthesis. The symbiotic bacteria living within the bodies of the larger invertebrate animals have been demonstrated to oxidize hydrogen sulfide, and the energy released from this biochemical process is used to power the fixation of carbon dioxide into small organic compounds – just as in free-living bacterial sulfide oxidation. The Calvin–Benson cycle employed in both cases is the same metabolic pathway that is utilized by plants in photosynthesis to transform inorganic carbon dioxide into organic compounds that are then utilized as food higher up in the food chain. The critical difference with chemosynthetic metabolism is that, rather than using sunlight, these animals and bacteria utilize chemical energy to power that reaction and are completely independent of sunlight (Figure 1). The net result is that freeliving bacteria in the environment and symbiotic bacteria living within animal tissues are able to live independently of sunlight by utilizing chemicals from the core of the earth, thus forming a very different basis for the food chain in the hydrothermal vent environment. The discovery of the hydrothermal vent environment was a fundamental discovery of a well-defined ecosystem that is completely independent of sunlight at any level of the food chain.
Figure 1 Chemosynthetic pathways in Riftia pachyptila.
The Ecophysiology of the Giant Tubeworm Riftia pachyptila One of the most dramatic and best-known of the animals endemic to the hydrothermal vent environment is the giant tubeworm Riftia pachyptila. Colonies of these worms are clumped together around effluent points in the hydrothermal vent habitat, growing toward and into the water that is percolating out from the seafloor. An individual animal lives inside a single, unbranched chitinous tube and the red structure protruding out of the end of the tube is the respiratory plume. The animal can retract the plume back into the tube if disturbed by a roaming predator. There is a collarlike vestimentum organ that positions the animal within the tube, and a large trunk region of the animal is filled with an organ termed the trophosome. This organ is believed to be the vestigial gut of the worm and is composed, literally, of masses of bacteria. A pool of coelomic fluid bathes the trophosome and contains a largemolecular-weight extracellular respiratory hemoglobin (Figure 2). Riftia pachyptila also has a separate pool of blood that contains high concentrations of an extracellular hemoglobin that circulates in an elaborate closed circulatory system powered by a heartlike structure in the vestimentum region. The blood is pumped in a complete circuit
Figure 2 The external anatomy of Riftia pachyptila.
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from the respiratory plume to body tissues, and on to elaborate capillary beds in the region of the trophosome and bacteria (Figure 3). The red color of the blood is due to the high concentration of hemoglobin and gives the characteristic red color to the plumes. The respiratory plume and the circulating hemoglobin are essential for the transport of the key metabolites oxygen, hydrogen sulfide, and carbon dioxide, which are the principal components of the metabolism of the symbiotic bacteria. The respiratory hemoglobins present in the plume and the coelomic fluid of the animal bind oxygen with a very high affinity. The binding is reversible and cooperative, such that oxygen uptake is enhanced at the respiratory plume, and oxygen delivery is augmented at the tissues and trophosome organ. Hydrogen sulfide is a highly toxic molecule that typically acts in a similar manner to cyanide by binding at the iron center of cytochrome molecules and hemoglobin molecules, thus arresting aerobic metabolism. Although Riftia pachyptila and other hydrothermal vent animals utilize hydrogen sulfide for their metabolism, they also have tissues that are highly sensitive to sulfide poisoning. Detoxification of hydrogen sulfide is essential for aerobic life in this
dynamic, chemically enriched environment. The key to the simultaneous needs for transportation and detoxification is the respiratory hemoglobin present in the plume and coelomic fluid. Riftia pachyptila hemoglobin binds hydrogen sulfide with a very high affinity. The toxic hydrogen sulfide is transported to the trophosome region in the center of the worm’s body as a tightly bound molecule that cannot chemically interact with sulfide-sensitive tissues. Oxygen and sulfide are simultaneously bound to the hemoglobin at separate binding sites and are transported to the trophosome, where they are believed to be delivered to the symbiotic bacteria for metabolism. In this way hydrogen sulfide is taken up from the surrounding sea water and transported to the site of bacterial metabolism while interaction is prevented with other tissues, such as the body wall, that are highly aerobic and sensitive to the toxic effects of hydrogen sulfide. These unusual adaptations function for respiratory gas transport and metabolism as well as for detoxification and tolerance of toxic chemicals in what would be a very inhospitable environment for most animals. Dense colonies of Riftia pachyptila flourish in a specialized microhabitat within the vent environment. The worms anchor themselves on the rocks where the hydrothermal vent fluid is issuing out into the seafloor. The base of the tube is bathed in hydrothermal fluid enriched in hydrogen sulfide and carbon dioxide, but devoid of oxygen. Temperatures are relatively elevated here, and a gradient develops along the length of the tube. The respiratory plume extends into ocean-bottom sea water that is 21C in temperature, devoid of hydrogen sulfide and enriched in oxygen (Figure 4). By occupying the interface between the hydrothermal fluids and the surrounding bottom water, animals are exposed to both of these essential metabolites which are then taken up by the circulating hemoglobin and
Figure 3 The internal anatomy of Riftia pachyptila.
Figure 4 The microhabitat of Riftia pachyptila.
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transported to internal tissues and symbiotic chemosynthetic bacteria.
The Ecophysiology of the Giant Clam Calyptogena magnifica The vesicomyid clam Calyptogena magnifica is a common inhabitant of hydrothermal vents that orients in the fissures of the pillow lava near the periphery of the vent environment. Individual clams can reach up to 20 cm in length. The clams position themselves in cracks in the pillow lava and wedge their muscular foot down into the region where the hydrothermal plume is percolating out (Figure 5). This water is enriched in hydrogen sulfide and is elevated in temperature. This position orients the siphon end-up into surrounding sea water and enables the uptake of oxygen and carbon dioxide. Calyptogena magnifica possess a high concentration of an intracellular, circulating hemoglobin that functions for oxygen binding and transport. The clam’s blood also contains a separate component that binds hydrogen sulfide and transports this essential metabolite to internal symbionts in the gill region. In this manner, the clams are able to accumulate hydrogen sulfide from the hydrothermal fluids bathing the foot, and oxygen via binding to the circulating hemoglobin through gill ventilation at the siphon region. These essential metabolites are transported as bound substances to symbiotic chemosynthetic bacteria in the gill via the circulatory system of the clam, providing both essential respiratory gas transport and detoxification of toxic hydrogen sulfide. The vent clams, like the tubeworms, seek out, exploit, and flourish in a unique microhabitat in the hydrothermal vent community. Other hydrothermal vent animals include the mytilid mussel Bathymodiolus thermophilus that also harbors chemosynthetic bacteria in its gill tissues. There are many free-ranging animals that are
Figure 5 The microhabitat of Calyptogena magnifica.
not fueled by chemosynthesis but may feed on the larger invertebrates that benefit from that symbiotic, chemoautotrophic metabolism. Animals such as the brachyuran crabs, Bythograea thermydron, that wander through the environment scavenging dead or dying material, numerous swarming amphipods, as well as slow-moving fishes, are plentiful. Along with the giant tubeworms, clams, and mussels, these animals benefit directly or indirectly from the chemicalbased metabolism that supports this dynamic and robust deep-sea community. It is the specialized physiological adaptations for transport and detoxification of hydrogen sulfide and other processes essential for life that provide the underlying mechanisms that make this possible.
Summary A fascinating variety of marine invertebrates occur in dense assemblages in organically enriched deep-sea hydrothermal vent environments. Intensive studies on the hydrothermal vent fauna have been conducted since their discovery in the late 1970s. Many investigations have focused on the fact that these organisms, including vestimentiferan tubeworms and vesicomyid clams emphasized here, are nutritionally dependent upon the chemical-based metabolism of large populations of symbiotic bacteria that they harbor internally in dense concentrations. These bacteria utilize hydrogen sulfide as an energy source to fix inorganic carbon into nutrients. Hydrogen sulfide is extremely toxic to aerobic organisms in nanomolar to micromolar concentrations. However, uptake and transport of sulfide to internal symbionts is essential for the host animal’s metabolism and survival. Chemical-based metabolism, or chemoautotrophy, and detoxification of sulfide through binding to blood-borne components occur in vent tubeworms and clams, and are particularly well-characterized for the tubeworm Riftia pachyptila. These chemosynthetic endosymbiont-harboring worms simultaneously transport sulfide bound to the respiratory hemoglobin, providing an electron donor for the bacterial symbiont metabolism and protection against sulfide toxicity at the tissues. How animals living in sulfide-rich environments like hydrothermal vent communities transport, metabolize, and detoxify hydrogen sulfide has been one of the major questions posed by hydrothermal vent researchers. The clarification of both the phylogenetic importance and the ecological status of these animals requires knowledge of their physiological adaptations to low oxygen conditions and high
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concentrations of sulfide, and provides answers to the puzzle of how these animals flourish in such seemly hostile environments.
See also Deep-Sea Ridges, Microbiology. Hydrothermal Vent Fluids, Chemistry of. Hydrothermal Vent Biota. Hydrothermal Vent Deposits. Hydrothermal Vent Ecology.
Further Reading Arp AJ, Childress JJ, and Fisher CR (1985) Blood gas transport in Riftia pachyptila. In: Jones ML (ed.) The Hydrothermal Vents of the Eastern Pacific: An Overview, Bulletin of the Biological Society of Washington, No. 6, p. 289. Washington DC: Biological Society of Washington. Childress JJ and Fisher CR (1992) The biology of hydrothermal vent animals: physiology, biochemistry and autotrophic symbioses. Oceanography and Marine Biology Annual Review 30: 337. Delaney JR (1998) Life on the seafloor and elsewhere in the solar system. Oceanus 41(2): 10. Feldman RA, Shank TB, Black MB, et al. (1998) Vestimentiferan on a whale fall. Biology Bulletin 194: 116.
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Fisher CR (1990) Chemoautotrophic and methanotrophic symbioses in marine invertebrates. Review of Aquatic Science 2: 399. Lutz RA (2000) Deep sea vents: science at the extreme. National Geographic (October): 116. MacDonald IR and Fisher CR (1996) Life without light. National Geographic (October): 313. Macdonald KC (1998) Exploring the global mid-ocean ridge. Oceanus 41(1): 2--9. Mullineaux L and Manahan D (1998) The LARVE Project explores how species migrate from vent to vent. Oceanus 41(2) Nelson K and Fisher CR (2000) Speciation of the bacterial symbionts of deep-sea vestimentiferan tube worms. Symbiosis 28: 1--15. Somero GN, Childress JJ, and Anderson AE (1989) Transport, metabolism, and detoxification of hydrogen sulfide in animals from sulfide-rich environments. Review of Aquatic Science 1: 591--614. Shank T, Fornari DJ, Von Damm KL et al. (1998) Temporal and spatial patterns of biological community development at nascent deep-sea hydrothermal vents (91500 N, East Pacific Rise). Deep-Sea Research II 45: 465. Tivey MK (1998) How to build a smoker chimney. Oceanus 41(2): 22. Van Dover CL (2000) The Ecology of Deep-Sea Hydrothermal Vents. Princeton, NJ: Princeton University Press.
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HYDROTHERMAL VENT FLUIDS, CHEMISTRY OF K. L. Von Damm, University of New Hampshire, Durham, NH, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 1246–1253, & 2001, Elsevier Ltd.
Introduction It was not until 1977 that we knew that fluids exit from the seafloor along the global midocean ridge system. With the first discovery of hydrothermal venting at the Galapagos Spreading Center, our ideas on how elements cycle through the oceans and lithosphere, and even how and where life on our planet may have originated were fundamentally and irrevocably changed. Although these fluids were only a few tens of degrees hotter than ambient sea water (o301C vs. 21C), on the basis of their chemical compositions it was immediately clear that these fluids were derived from reactions at much higher temperatures between sea water and the oceanic crust. Less than two years later, spectacular jets of hot (X3501C) and black water were discovered several thousand kilometers away on the northern East Pacific Rise, and ‘black smokers’ and ‘chimneys’ (Figure 1) entered the oceanographic lexicon. From a chemical oceanography perspective, hydrothermal venting and hydrothermal vent fluids provide both new source and sink mechanisms for elemental cycling in the ocean, and therefore possible resolutions to a number of the outstanding chemical flux imbalances. They therefore play a fundamental role in regulating the chemistry of the oceans through geological time. From a biological oceanography perspective, hydrothermal vents provide us with new ecosystems based on chemosynthetic, rather than photosynthetic energy. Of the 4500 new species discovered at these sites, the archea and other microbiological components are attracting increasing interest for both biotechnological applications and ‘origin of life’ questions on our own and other planets. From a physical oceanographic perspective, hydrothermal vents provide an input of both heat and materials into the oceanic mid-depth circulation. From a geological oceanographic perspective they provide an efficient means of removing heat from newly formed oceanic crust, as well as a means of altering the elements recycled into the mantle when the oceanic crust formed at spreading centers is later subducted back into the Earth’s interior.
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Where Are Hydrothermal Vents Found? Hydrothermal vents are now known to exist at approximately 30 locations on the global midocean ridge system (Figure 2). Initially some people had speculated that venting would only be found on intermediate- or faster-spreading ridges (i.e., ridges with full spreading rates of at least 60 mm yr1); we now know of numerous locations on slow-spreading ridges (e.g., Mid-Atlantic Ridge) where they occur. Known sites occur at depths from 800 to 43600 m, with spreading rates from o20 to 4150 mm y1 (full rate), on both bare basalt and sedimented-covered ridges, as well as on seafloor where ultramafic rock types are known to outcrop, and at temperatures up to 4051C. If one looks at the global distribution of known vent sites, it is obvious that many of the sites are in relatively close proximity to nations that operate submersibles, and are clustered disproportionately in the north Atlantic and eastern Pacific. Although not strictly part of the midocean ridge system, venting associated with back arc spreading centers is also known from a number of sites in the western Pacific. No sites have yet been discovered in the Indian Ocean, although cruises to this ocean are now planned. Similarly, no sites are known in the south Atlantic, or at high latitudes. This is an exploration issue, not a lack of their existence in these areas. While initially vents were thought to occur at the mid-point of ridge segments, this was a largely selffulfilling prophesy, as this is where exploration for them was focused. There is increasing evidence that more venting occurs on the magmatically robust portions of the ridge, rather than on those areas that are deemed to be magma-starved, on the basis of their morphological characteristics.
How Are They Found? Vent fields have been discovered in numerous ways, but surveys of the overlying water column and camera tows are the most common systematic approaches employed today. The venting of hot/warm water often forms a plume with unique temperature, salinity, reduced light transmittance, and other specific chemical signals several hundred meters above the ridge. The presence of such a plume is often the first indication that a given section of ridge is hydrothermally active. Cameras, or other
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(B)
(A)
(C)
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(
D)
(E) Figure 1 Pictures of black smokers. (A)–(C) Brandon Vent a black smoker with 4051C fluid temperatures on the Southern East Pacific Rise at 211330 S at a depth of 2834 m; in (A) and (B), shown instrumented with a Hobo recording temperature probe and in (C) with one of Alvin’s manipulators. (D) Nadir vent at 171260 S at a depth of 2562 m and fluid temperatures of 3431C shown prior to sampling in 1998. In the foreground is Alvin’s basket with water sampling equipment. (E) Water sampling of Nadir vent fluids.
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Figure 2 Global distribution of known (sampled) hydrothermal systems. (Modified from Baker et al. (1995).)
deep-towed survey vehicles may also find visual evidence for communities of vent-specific organisms, sulfide structures (both active and extinct), and murky water. These surface ship surveys may then lead to the use of an ROV (remotely operated vehicle) or submersible, in what has been called a ‘nested survey strategy.’ Sites have also been found by the fortuitous dredging up of both pieces of hydrothermal chimneys and vent-specific animals.
Controls on Fluid Compositions Hydrothermal vent fluids are primarily, if not entirely, sea water that has been altered due to reaction
within the oceanic crust at high temperatures (at least 3501C, and more likely 44001C) and elevated pressures (at least 80 bar, and more typically at least 250 bar). The two primary controls on vent fluid compositions are phase separation and water–rock reaction (Figure 3). Unlike pure water, sea water is a two-component system, containing both water and salt (primarily NaCl, especially for hydrothermal fluids from which the magnesium and sulfate have been essentially quantitatively removed). The critical point for sea water is higher than that for distilled water – 4071C and 298 bar rather than 3741C and 220 bar – and most importantly the two-phase curve therefore continues beyond the critical point
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Sea water Oceanic crust 4
CaSo 2+
Mg
W_R rxn 2+
+
Ca , H W_R rxn n
io
t ra pa
Vapor e
s ha
Brine
se
P
W_R rxn
Gas input Heat source Magma lens or dike
Figure 3 Schematic of a hydrothermal flow cell. W– Rrxn ¼ water–rock reaction. (Von Damm 1995 in Humphris et al.)
(Figure 4). If fluids intersect the two-phase curve at pressure and temperature conditions less than the critical point, they will undergo subcritical phase separation or boiling. The vapor phase formed by
0 50
Vapor + halite
100
Pressure (bar)
150 200 250 300
Liquid
CP Liquid + vapor
350
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this process will contain some salt, the exact amount and composition of which will depend on the pressure and temperature conditions at which the phase separation occurred, and will be enriched in dissolved gases. If the fluid intersects the two phase curve at pressure and temperature conditions higher than the critical point, supercritical phase separation or condensation will occur, and, rather than a lowsalinity vapor being formed, a high-salinity brine or liquid phase will condense. In most, if not all, hydrothermal fluids that have been sampled, phase separation has occurred as evidenced by the chlorinities of these fluids, which can be from B6% to 200% of those in sea water. While the change in the absolute chlorinity of seafloor hydrothermal fluids is primarily the result of phase separation, changes in the other dissolved ions in sea water also reflect substantial reaction of the fluids with the rock, mostly basalt, but also ultramafics and sometimes sediments. Seafloor hydrothermal fluids therefore do not show the ‘constancy of composition’ and elemental ratios known to characterize the major chemical species in sea water. The two primary mechanisms for determining the composition of hydrothermal fluids are therefore phase separation and water–rock interaction. This does not mean that other mechanisms may not be important. Two that have been identified but whose global importance is not yet known are magmatic degassing and biological uptake/removal. On the East Pacific Rise at 91500 N latitude unusually high gas concentrations and gas/heat ratios have been observed in the hydrothermal fluids that suggest that we are seeing the degassing of a recently resupplied magma chamber. This has not yet been observed at any other sites, although it has been observed for B7 years at this site. The potential biological effects on fluid compositions have long been speculated on, but there are few quantitative data at this time that can directly address this question. As the ‘black smoker’ fluids are at temperatures well beyond the known bounds of life on this planet, if biological effects are present they are most likely to be found in fluids with temperatures less than B1101C.
400
The Division of Fluids by Temperature and Styles of Venting
450 500 250
300
350
400
450
500
Temperature (°C) Figure 4 Two-phase curve for sea water, showing the location of the critical point (CP) for sea water (4071C, 298 bar), and the relative locations of the liquid, liquid þ vapor, and vapor þ halite stability fields. Note that, unlike for pure water, the two-phase curve continues beyond the critical point.
Hydrothermal fluids are often subdivided into two categories: high-temperature or focused flow, and lowtemperature or diffuse flow. These terms are often poorly defined in the contexts in which they are used, and may mean different things to different authors. High-temperature fluids, generally 4 200–2501C, are usually focused jets of water that are exiting from
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constructional features that either are precipitated from the hydrothermal fluids themselves when they encounter cold, alkaline sea water or are formed by precipitation due to the mixing of hydrothermal fluids with sea water, or, and more usually, both of these processes. Although these chimneys comprise metal sulfides and sulfates, other minerals are also present, usually in lesser quantities, although on sedimentcovered ridges carbonates, for example, may also be abundant in these chimney or mound structures. Lowtemperature fluids, generally p351C, but sometimes approaching 1001C, have usually deposited (if they ever contained) much of their metal load below the seafloor and hence usually neither smoke nor have specific constructional features associated with their fluid flow. Because of this lack of specific structure, the fluid flow is often less organized, and hence is referred to as ‘diffuse.’ Some authors use this term to refer to fluids exiting directly from the basaltic substrate. These sites are usually well-colonized by various types of vent megafauna, making it difficult, if not impossible, to identify a specific orifice. Other authors use this same term to refer to fluids that are oozing out of large sulfide structures. At some sites, only hightemperature ‘black smoker’ vents are found, while at other sites only the lower-temperature diffuse venting is found (as at the Galapagos Spreading Center where venting was first discovered). At other sites, both focused and diffuse flow may occur right next to one another, and, at least at one site, we have now observed low-temperature diffuse flow evolving into high-temperature focused flow over several years. For all types of ‘low temperature and/or diffuse flow,’ all of the fluids that have been sampled to date appear to contain some fluid that has reacted at significantly greater temperatures than are measured directly in these fluids. Hence one could argue that, at least at the ridge axes, all fluids are ultimately high-temperature fluids and that some of them have undergone significant mixing with sea water at some depth, likely relatively shallow, within the oceanic crust. With the continued inability to drill these systems, our knowledge of their subsurface hydrology is rudimentary at best. One might note that fluids with temperatures in the B100–2001C range have been left out of the above classification. This is because there are few fluids that have been sampled in this temperature range, which may reflect a true lack of abundance of fluids at such temperatures, or simply a sampling gap.
were chemically stable for years at a time, yet each individual vent had a distinctive chemistry, and rapid changes in the hydrothermal plumes above these sites had been noted. In 1991 some of this puzzle was resolved with the first opportunity to sample a midocean ridge hydrothermal site immediately following a volcanic eruption and to study the evolution of this site on timescales now approaching decadal. The discovery of an eruption on the East Pacific Rise at 9145–520 N was fortuitous, but it could be sampled and documented with the DSV Alvin within weeks of the volcanic event. Since that time using the US Navy’s SOSUS system, it has been possible to determine real-time when some of these events are occurring on the Juan de Fuca and Gorda Ridges, but immediate response to these events have been limited to surface ship observations due to logistical constraints. The events studied so far have been characterized by a volcanic eruption at the seafloor, the intrusion of a dike, or both. The knowledge gained from responding to these events has revolutionized our understanding of these systems, especially their pronounced temporal variability on very short timescales (much less than days to weeks). Presumably we will be able to study more of these events and to gain a sense of their frequency, perhaps as a function of spreading rate, as we begin to instrument more of the ridge crest with hydrophones to detect the T-phase signals associated with these events. The Influence of Tectonic/Cracking Events
Magmatic events have provided new insights into the processes that drive hydrothermal systems and their fluid compositions on intermediate- to fast-spreading ridges, but presumably tectonic events are more important (or at least more frequent than magmatic events) on slow-spreading ridges. As we have not yet been able to observe one of these events, we cannot assess their importance. In a relatively small cracking event on a fast-spreading ridge observed with a seismic array, changes in fluid temperatures were marked, and changes in fluid compositions were profound, leading to major changes in the biological communities existing at this site. Presumably the observation of one or more tectonic events on slowspreading ridges will also provide critical new insights into how hydrothermal systems on these types of ridges function and evolve. On-Axis versus Off-Axis
The Influence of Volcanic Events
In the late 1980s there was a real dichotomy in our understanding of the controls on seafloor hydrothermal systems. We knew of individual vents that
The discussion above has focused on the axial component of midocean ridge hydrothermal systems. Many debates have focused on the importance of these axial systems compared to hydrothermal
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HYDROTHERMAL VENT FLUIDS, CHEMISTRY OF
systems on the ridge flanks. The net amount of heat transported and the total geographical extent of the flanks is much larger than that of the axial parts of the ridges, the division often being drawn at 1 My, resulting in a partitioning of 70%:30% relative heat loss. Hence, the relative chemical anomaly and volume of fluids expelled on the axis versus the flanks of the ridge will determine which of these parts of the midocean ridge hydrothermal system is most important for the cycles of individual chemical elements and species. It has been difficult to acquire quantitative data on the fluids exiting from the flanks, one of the reasons being that sites of flank fluid venting are difficult to identify because water column plumes are not present, nor are the distinctive animal communities (distinctive both for their animal types as well as their white color on the black basalt) found at the axial sites of venting. In the last several years drilling by the Ocean Drilling Program on the flanks of the Juan de Fuca Ridge has provided important new information on the chemistry and hydrology of these systems, as well as demonstrating a viable method for further approaching the flank question.
Table 1
Observed Fluid Compositions Compositions of hydrothermal fluids not only vary widely but also are almost always very different from those of sea water (Table 1). While some of the low temperature diffusely venting fluids may be close to sea water in their major element compositions, they will often have very different compositions of dissolved gases (e.g., H2S, CO2, CH4, H2, and He) and will usually be highly enriched in iron and manganese compared to local ambient sea water. Compared to sea water, hydrothermal fluids have lost essentially all of their magnesium and sulfate, and are highly enriched in H2S, CH4, H2, He, Si, Li, Fe, and Mn. As hydrothermal fluids are very acid, they also have no alkalinity, and with the loss of sulfate, chloride becomes the major, and almost only anion (bromide is present in much lower concentrations). The behavior of the cations is more variable. As the amount of chloride present is a result of the phase separation history of the fluids, and the fluids must maintain electroneutrality, to determine whether a particular cation has been added to or removed from the fluid,
Range of physical parameters and chemical compositions for hydrothermal vent fluidsa
Parameter
Unitsb
Vents
Seawater
Temperature Depth (pressure) pH 251C, 1 atm Alkalinitytotal Cl SO24 H2S Si Li Na K Ca Mg Sr Fe Mn Cu Zn Cd Co Pb B Al Br CO2 CH4 H2
1C m (bars)
4 2–405 800–3600 (80–360) 2.5–7.8 2.7–10.6 31–1245 0 0–110 2.7b–24.0 o0.012–2.35 o15–924 o1–58.7 o0.2–109 0 o1–348 0–18.7 0–4.48 0–310 0–900 o10–1000 o5–2570 50–2200 416–1630 0–20.0 29–1880 2.3–375 0.0003–6800 o0.001–38
B1–5 (bottom water)
a b
meq kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 nmol kg1 nmol kg1 nmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1 mmol kg1
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Only includes mid-ocean ridge systems, both bare rock and sediment-covered. From high-temperature vents.
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7.8 2.3 545 28 0 0.032–0.180 0.026 465 10 10 53 87 o0.001 o0.001 0.007 0.01 1 0.03 0.01 416 0.02 840 2.3 0.0003 0.0000003
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not only the absolute concentration but also its ratio to chloride must be compared to the local ambient sea water value. Hydrothermal fluids are also very reducing, often containing large amounts of H2S, H2, and CH4. In most, but not all hydrothermal fluids (fluids that are formed immediately after a volcanic eruption being a major exception), Li, K, Rb, Cs, Ca, Sr, Si, the transition metals in reduced forms, including Fe, Mn, Cu, and Zn, are enriched in hydrothermal fluids, the cause being water–rock interaction. Sodium may be either enriched or depleted with respect to the chloride content of the fluids, the loss being due to albitization, with a concomitant gain in the calcium-content of the fluids. It is the loss of magnesium to form magnesiumhydroxy silicates that, along with other aluminosilicate reactions, generates and then maintains the acidity of the fluids. In the case of sulfate, some is lost as anhydrite (CaSO4) in the downflow zone, while some is reduced to sulfide (as H2S). The compositions of fluids exiting from a single hydrothermal vent may vary widely over time. The cases in which this has been observed are increasing, and are usually associated with vents where a known magmatic event has occurred. The variation in the composition of a single vent can vary from vapor to brine, and may encompass almost the entire range of known compositions. In contrast, some sites of venting are known where the fluid compositions have been stable during the time interval over which they have been sampled. None of these vents with constant compositions have known ‘magmatic’ or ‘tectonic’ events associated with them, although several of these sites have now been sampled over times of B15 years. Presumably, these vents are in a period of steady-state venting, although we do not have adequate constraints to determine how long after an eruptive event, or at what spreading rates, this may occur. While our data has increased on the temperature and chemical characteristics of vents during their early histories, few data exist for their waning stage(s). Presumably all this variability – or lack thereof – and the timescale(s) on which it occurs can ultimately be tied to the nature of the heat source at a given site. Aside from the most general characteristics related to the presence or absence of a seismic low-velocity zone, and the depth at which it occurs, little is known about the specifics of the heat sources at sites that are hydrothermally active, especially in contrast to those that are not. The composition of a hydrothermal fluid cannot be correlated to, or predicted by, such known physical parameters as the depth of the seafloor on which it occurs, the spreading rate of the ridge on which it occurs, and so on. As the fluid compositions in most
cases are probably due to the achievement of steadystate, if not true thermodynamic equilibrium, of the fluids with the rock substrate, some of the measured compositions can be tied to either the measured exit temperature or the presumed in situ conditions within the hydrothermal system itself. While there has been some success, especially recently, with understanding the chemical controls on these systems using thermodynamic modeling, a major limitation in many cases remains the proximity to the critical point of both pure water and sea water.
The Flux Question One of the driving questions for the study of seafloor hydrothermal systems is to understand their net flux to the ocean in terms of energy (thermal and chemical) and mass. The thermal, or heat, energy they carry is believed to be relatively well constrained, as various independent ways of estimating this flux provide similar values. The mass of chemicals they add and/or remove remains problematic. In some cases whether hydrothermal activity is a net source or sink for particular elements remains unresolved as well. In addition to absolute concentrations, or concentrations normalized to the chloride content, the isotopic signature of various species can also be used to constrain the source and sink terms, as well as helping to identify the important processes occurring within the hydrothermal circulation cell.
Summary and Conclusions Hydrothermal venting along the global midocean ridge system is a process that is widespread throughout the ocean basins and impacts all of the oceanographic disciplines. Our studies of these systems remain in their infancy, however, and we do not yet completely understand the controls on the chemistry of these systems, the controls on the locations of individual vent sites, their overall importance to ocean chemistry, productivity, and circulation, and their net effects on the structure and composition of the oceanic crust.
See also Hydrothermal Vent Deposits. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. MidOcean Ridge Geochemistry and Petrology. MidOcean Ridge Seismic Structure. Seamounts and Off-Ridge Volcanism.
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Further Reading Cowen JP and Baker ET (1998) Topical studies in oceanography: detection of and response to mid-ocean ridge magmatic events. Deep-Sea Research II 45(12): 2503--2766. Elderfield H and Schultz A (1996) Mid-ocean ridge hydrothermal fluxes and the chemical composition of the ocean. Annual Review of Earth and Planetary Science 24: 191--224. Humphris SE, Zierenberg RA, Mullineaux LS, and Thomson RE (eds.) (1995) Seafloor Hydrothermal Systems: Physical, Chemical, Biological, and Geological
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Interactions. AGU Monograph 91. Washington, DC: American Geophysical Union. Parson LM, Walker CL, and Dixon DR (eds.) (1995) Hydrothermal Vents and Processes. London: Geological Society: Geological Society Special Publication 87. Seyfried WE Jr (1987) Experimental and theoretical constraints on hydrothermal alteration processes at mid-ocean ridges. Annual Review of Earth and Planetary Science 15: 317--335. Von Damm KL (1990) Seafloor hydrothermal activity: black smoker chemistry and chimneys. Annual Review of Earth and Planetary Science 18: 173--204.
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HYPOXIA N. N. Rabalais, Louisiana Universities Marine Consortium, Chauvin, LA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction: Definitions Hypoxic (low-oxygen) and anoxic (no-oxygen) waters have existed throughout geologic time. Presently, hypoxia occurs in many of the ocean’s deeper environs, open-ocean oxygen-minimum zones (OMZs), enclosed seas and basins, below western boundary current upwelling zones, and in fjord. Hypoxia also occurs in shallow coastal seas and estuaries, where their occurrence and severity appear to be increasing, most likely accelerated by human activities (Figure 1). A familiar term used in the popular press and literature, ‘dead zone’, used for coastal and estuarine hypoxia, refers to the fish and shellfish killed by the suffocating conditions or the failure to catch these animals in bottom waters when the oxygen concentration in the water covering the seabed is below a critical level.
Based on laboratory or field observations or both, the level of oxygen stress and related responses of invertebrate and fish faunas vary. The units are often determined by oxygen conditions that are physiologically stressful, but these levels also differ depending on the organisms considered, and the pressure, temperature, and salinity of the ambient waters. The numerical definition of hypoxia varies as do the units used, but hypoxia has mostly been defined as dissolved oxygen levels lower than a range of 3–2 ml l 1, with the consensus being in favor of 1.4 ml l 1 ( ¼ 2 mg l 1 or ppm). This value is approximately equivalent to 30% oxygen saturation at 25 1C and 35 salinity (psu). Below this concentration, bottom-dragging trawl nets fail to capture fish, shrimp, and swimming crabs. Other fishes, such as rays and sharks, are affected by oxygen levels below 3 mg l 1, which prompts a behavioral response to evacuate the area, up into the water column and shoreward. Water-quality standards in the coastal waters of Long Island Sound, New York, and Connecticut, USA, consider that dissolved oxygen conditions below 5 mg l 1 result in behavioral effects in marine organisms and fail to support living resources at sustainable levels.
Oxygen depletion Annual Episodic Periodic Persistent Unknown
Figure 1 Distribution of coastal ocean hypoxic areas; excludes naturally occurring hypoxia, such as upwelling zones and OMZs. Reproduced from Dı´az RJ, Nestlerode J, and Dı´az ML (2004) A global perspective on the effects of eutrophication and hypoxia on aquatic biota. In: Rupp GL and White MD (eds.) Proceedings of the 7th International Symposium on Fish Physiology, Toxicology and Water Quality, Tallinn, Estonia, May 12-15, 2003. EPA 600/R-04/049, pp. 1–33. Athens, GA: Ecosystems Research Division, US Environmental Protection Agency, with permission from Robert J. Dı´az.
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The most commonly used definition for oceanic waters is dissolved oxygen content less than 1 ml l 1 (or 0.7 mg l 1). Disoxyic or disaerobic refers to oxygen levels between 0.1 and 1.0 ml l 1. OMZs are usually defined as waters less than 0.5 ml l 1 dissolved oxygen.
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shifts in trophic interactions and food webs, and impacts on living resources.
Hypoxic Systems Oxygen-minimum Zones
Causes Hypoxia occurs where the consumption of oxygen through respiratory or chemical processes exceeds the rate of supply from oxygen production via photosynthesis, diffusion through the water column, advection, or mixing. The biological and physical water-column characteristics that support the development and maintenance of hypoxia include (1) the production, flux, and accumulation of organic-rich matter from the upper water column; and (2) watercolumn stability resulting from stratification or long residence time. Dead and senescent algae, zooplankton fecal pellets, and marine aggregates contribute significant amounts of organic detritus to the lower water column and seabed. Aerobic bacteria consume oxygen during the decay of the carbon and deplete the oxygen, particularly when stratification prevents diffusion of oxygen. Stratification is the division of the water column into layers with different densities caused by differences in temperature or salinity or both. Hypoxia will persist as long as oxygen consumption rates exceed those of supply. Oxygen depletion occurs more frequently in estuaries or coastal areas with longer water residence times, with higher nutrient loads and with stratified water columns. Hypoxia is a natural feature of many oceanic waters, such as OMZs and enclosed seas, or forms in coastal waters as a result of the decomposition of high carbon loading stimulated by upwelled nutrientrich waters. Hypoxia in many coastal and estuarine waters, however, is but one of the symptoms of eutrophication, an increase in the rate of production and accumulation of carbon in aquatic systems. Eutrophication very often results from an increase in nutrient loading, particularly by forms of nitrogen and phosphorus. Nutrient over-enrichment from anthropogenic sources is one of the major stressors impacting estuarine and coastal ecosystems, and there is increasing concern in many areas around the world that an oversupply of nutrients is having pervasive ecological effects on shallow coastal waters. These effects include reduced light penetration, increased abundance of nuisance macroalgae, loss of aquatic habitat such as seagrass or macroalgal beds, noxious and toxic algal blooms, hypoxia and anoxia,
Persistent hypoxia is evident in mid-water OMZs, which are widespread in the world oceans where the oxygen concentrations are less than 0.5 ml l 1 (or about 7.5% oxygen saturation, o22 mM). They occur at different depths from the continental shelf to upper bathyal zones (down to 1300 m). Many of the OMZs form as a result of high primary production associated with coastal upwelled nutrient-rich waters. Their formation also requires stagnant circulation, long residence times, and the presence of oxygen-depleted source waters. The extensive OMZ development in the eastern Pacific Ocean is attributed to the fact that intermediate depth waters of the region are older and have overall oxygen concentrations lower than other water masses. The largest OMZs are at bathyal depths in the eastern Pacific Ocean, the Arabian Sea, the Bay of Bengal, and off southwest Africa. The upper boundary of an OMZ may come to within 10 or 50 m of the sea surface off Central America, Peru, and Chile. The OMZ is more than 1000-m thick off Mexico and in the Arabian Sea, but off Chile, the OMZ is o400-m thick. Along continental margins, minimum oxygen concentrations occur typically between 200 and 700 m. The area of the ocean floor where oceanic waters permanently less than 0.5 ml l 1 impinge on continental margins covers 106 km2 of shelf and bathyal seafloor, with over half occurring in the northern Indian Ocean. These permanently hypoxic waters account for 2.3% of the ocean’s continental margin. These hypoxic areas are not related to eutrophication, but longer-term shifts in meteorological conditions and ocean currents may increase their prevalence in the future with global climate change. Shifts in ocean currents have been implicated in the increased frequency of continental shelf hypoxia along the northwestern US Pacific coast of Oregon. Deep Basins, Enclosed Seas, and Fjord
Many of the existing permanent or periodic anoxic ocean environments occur in enclosed or semienclosed waters where a mass of deep water is bathymetrically isolated from main shelf or oceanic water masses by surrounding landmasses or one or more shallow sills. In conjunction with a pycnocline, the bottom water volume is restricted from exchange with deep open water. Examples of hypoxic and
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anoxic basins include anoxic deep water fjord, such as Saanich Inlet, the deeper basins of the Baltic, the basin of the Black Sea, the Japanese Seto Inland Sea, deep waters of the Sea of Cortez, Baja California, and Santa Barbara Basin in the southern California borderland. Coastal Seas and Estuaries
Periodic hypoxia or anoxia also occurs on open continental shelves, for example, the northern Gulf of Mexico and the Namibian and Peruvian shelves where upwelling occurs. More enclosed shelves such as the northern Adriatic and the northwestern shelf of the Black Sea also have periodic hypoxia or anoxia. In these instances, there is minimal exchange of shelf-slope water and/or high oxygen demand on the shallow shelf. Estuaries, embayments, and lagoons are susceptible to the formation of hypoxia and anoxia if the water residence time is sufficiently long, especially where the water column is stratified. Light conditions are also important in these coastal habitats as a limiting factor on phytoplankton growth, which, if excessive, contributes to high organic loading within the confined waters. Coastal ecosystems that have been substantially changed as a result of eutrophication exhibit a series of identifiable symptoms, such as reduced water clarity, excessive, noxious, and, sometimes, harmful algal blooms, loss of critical macroalgal or seagrass habitat, development or intensification of oxygen depletion in the form of hypoxia or anoxia, and, in some cases, loss of fishery resources. More subtle responses of coastal ecosystems to eutrophication include shifts in phytoplankton and zooplankton communities, shifts in the food webs that they support, loss of biodiversity, changes in trophic interactions, and changes in ecosystem functions and biogeochemical processes. In a review of anthropogenic hypoxic zones in 1995, Dı´az and Rosenberg noted that no other environmental variable of such ecological importance to estuarine and coastal marine ecosystems around the world has changed so drastically, in such a short period of time, as dissolved oxygen. For those reviewed, there was a consistent trend of increasing severity (either in duration, intensity, or size) where hypoxia occurred historically, or hypoxia existed presently when it did not occur before. While hypoxic environments have existed through geologic time and are common features of the deep ocean or adjacent to areas of upwelling, their occurrence in estuarine and coastal areas is increasing, and the trend is consistent with the increase in human activities that result in nutrient over-enrichment.
The largest human-caused hypoxic zone is in the aggregated coastal areas of the Baltic Sea, reaching 84 000 km2. Hypoxia existed on the northwestern Black Sea shelf historically, but anoxic events became more frequent and widespread in the 1970s and 1980s, reaching over areas of the seafloor up to 40 000 km2 in depths of 8–40 m. There is also evidence that the suboxic zone of the open Black Sea enlarged toward the surface by about 10 m since 1970. The condition of the northwestern shelf of the Black Sea, in which hypoxia covered up to 40 000 km2, improved over the period 1990–2000 when nutrient loads from the Danube River decreased, but may be experiencing a worsening of hypoxic conditions more recently. Similar declines in bottom water dissolved oxygen have occurred elsewhere as a result of increasing nutrient loads and cultural eutrophication, for example, the northern Adriatic Sea, the Kattegat and Skaggerak, Chesapeake Bay, Albemarle-Pamlico Sound, Tampa Bay, Long Island Sound, New York Bight, the German Bight, and the North Sea. In the United States, over half of the estuaries experience hypoxia at some time over an annual period and many experience hypoxia over extensive areas for extended periods on a perennial basis. The number of estuaries with hypoxia or anoxia continues to rise. Historic data on Secchi disk depth in the northern Adriatic Sea in 1911 through the present, with few interruptions of data collection, provide a measure of water transparency that could be interpreted to depict surface water productivity. These data coupled with surface and bottom water dissolved oxygen content determined by Winkler titrations and nutrient loads outline the sequence of eutrophication in the northern Adriatic Sea. Similar historical data from other coastal areas around the world demonstrate a decrease in water clarity due to phytoplankton production in response to increased nutrient loads that are paralleled by declines in water column oxygen levels. There are strong relationships between river flow and nutrient flux into the Chesapeake Bay and northern Gulf of Mexico and phytoplankton production and biomass and the subsequent fate of that production in spring deposition of chlorophyll a. Further there is a strong relationship between the deposited chlorophyll a and the seasonal decline of deep-water dissolved oxygen. Excess nutrients in many watersheds are driven by agricultural activities and atmospheric deposition from burning of fossil fuels. The link with excess nutrients in more urban areas, such as Long Island Sound, is with the flux of nutrients associated from numerous wastewater outfalls.
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HYPOXIA
Swift currents that move materials away from a river delta and that do not permit the development of stratification are not conducive to the accumulation of biomass or depletion of oxygen, for example in the Amazon and Orinoco plumes. Similar processes off the Changjiang (Yantze River) and high turbidity in the plume of the Huanghe (Yellow River) were once thought to be reasons why hypoxia did not develop in those coastal systems. Incipient indications of the beginning of symptoms of cultural eutrophication were becoming evident at the terminus of both these systems as nutrient loads increased. The severely reduced, almost minimal, flow of the Huanghe has prevented the formation of hypoxia, but other coastal ecosystem problems remain. There is, however, now a hypoxic area off the Changjiang Estuary and harmful algal blooms are more frequent in the East China Sea. The likelihood that more and more coastal systems, especially in developing countries, where the physical conditions are appropriate will become eutrophic with accompanying hypoxia is worrisome. Northern Gulf of Mexico
The hypoxic zone on the continental shelf of the northern Gulf of Mexico is one of the largest hypoxic zones in the world’s coastal oceans, and is representative of hypoxia resulting from anthropogenic activities over the last half of the twentieth century (Figure 2). Every spring, the dissolved oxygen levels in the coastal waters of the northern Gulf of Mexico decline and result in a vast region of oxygen-starved water that stretches from the Mississippi River westward along the Louisiana shore and onto the Texas coast. The area of bottom covered by hypoxic water can reach 22 000 km2, and the volume of hypoxic waters may be as much as 1011 m3. Hypoxia
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in the Gulf of Mexico results from a combination of natural and human-influenced factors. The Mississippi River, one of the 10 largest in the world, drains 41% of the land area of the lower 48 states of the US and delivers fresh water, sediments, and nutrients to the Gulf of Mexico. The fresh water, when it enters the Gulf, floats over the denser saltier water, resulting in stratification, or a two-layered system. The stratification, driven primarily by salinity, begins in the spring, intensifies in the summer as surface waters warm and winds that normally mix the water subside, and dissipates in the fall with tropical storms or cold fronts. Hypoxic waters are found at shallow depths near the shore (4–5 m) to as deep as 60 m. The more typical depth distribution is between 5 and 35 m. The hypoxic water is not just located near the seabed, but may rise well up into the water column, often occupying the lower half of a 20-m water column (Figure 3). The inshore/offshore distribution of hypoxia on the Louisiana shelf is dictated by winds and currents. During typical winds from the southeast, downwelling favorable conditions force the hypoxic bottom waters farther offshore. When the wind comes from the north, an upwelling favorable current regime promotes the movement of the hypoxic bottom waters close to shore. When the hypoxic waters move onto the shore, fish, shrimp, and crabs are trapped along the beach, resulting sometimes in a ‘jubilee’ when the stunned animals are easily harvested by beachgoers. A more negative result is a massive fish kill of all the sea life trapped without sufficient oxygen. Hypoxia occurs on the Louisiana coast west of the Mississippi River delta from February through November, and nearly continuously from mid-May through mid-September. In March and April, hypoxic water masses are patchy and ephemeral. The hypoxic
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Figure 2 Similar size and expanse of bottom water hypoxia in mid-July 2002 (shaded area) and in mid-July 2001 (outlined with dashed line). Data source: N. N. Rabalais, Louisiana Universities Marine Consortium.
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Distance (km) Figure 3 Contours of dissolved oxygen (mg l 1) across the continental shelf of Louisiana approximately 200 km west of the Mississippi River delta in summer. The distribution across the shelf in August is a response to an upwelling favorable oceanographic regime and that of September to a downwelling favorable oceanographic regime. These contours also illustrate the height above the seabed that hypoxia can reach, i.e., over half the water column. Data source: N. N. Rabalais, Louisiana Universities Marine Consortium.
zone is most widespread, persistent, and severe in June, July, and August, and often well into September, depending on whether tropical storm activity begins to disrupt the stratification and hypoxia. Anoxic waters occur periodically in midsummer. The midsummer size of the hypoxic zone varies annually, and is most closely related to the nitrate load of the Mississippi River in the 2 months prior to the typically late-July mapping exercise. The load of nitrate is determined by the discharge of the Mississippi River multiplied by the concentration of the nitrate, so that the amount of water coming into the Gulf of Mexico is also a factor. The relationship of the size of hypoxia, however, is stronger with the load of nitrate than with the total river water discharge or any other nutrient or combination of
nutrients. Changes in the severity of hypoxia over time are related mostly to the change in nitrate concentration in the Mississippi River (80%), the remainder to changes in increased discharge (20%).
Historical Change in Oxygen Conditions Historical dissolved oxygen data such as those for the northern Adriatic Sea beginning in 1911 are not commonly available. A solution is to turn to the sediment record for paleoindicators of long-term transitions related to eutrophication and oxygen deficiency. Biological, mineral, or chemical indicators of plant communities, level of productivity, or
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HYPOXIA
conditions of hypoxia preserved in sediments, where sediments accumulate, provide clues to prior hydrographic and biological conditions. Data from sediment cores taken from the Louisiana Bight adjacent to the Mississippi River where sediments accumulate with their biological and chemical indicators document increased recent eutrophication and increased organic sedimentation in bottom waters, with the changes being more apparent in areas of chronic hypoxia and coincident with the increasing nitrogen loads from the Mississippi River system beginning in the 1950s. This evidence comes as an increased accumulation of diatom remains and marine-origin carbon accumulation in the sediments. Benthic microfauna and chemical conditions provide several surrogates for oxygen conditions. The mineral glauconite forms under reducing conditions in sediments, and its abundance is an indication of low-oxygen conditions. (Note that glauconite also forms in reducing sediments whose overlying waters are 42 mg l 1 dissolved oxygen.) The average glauconite abundance in the coarse fraction of sediments in the Louisiana Bight was B5.8% from 1900 to a transition period between 1940 and 1950, when it increased to B13.4%, suggesting that hypoxia ‘may’ have existed at some level before the 1940–50 time period, but that it worsened since then. Benthic foraminiferans and ostracods are also useful indicators of reduced oxygen levels because oxygen stress decreases their overall diversity as measured by the Shannon–Wiener diversity index (SWDI) and shifts community composition. Foraminiferan and ostracod diversity decreased since the 1940s and early 1950s, respectively. While presentday foraminiferan diversity is generally low in the Louisiana Bight, comparisons among assemblages from areas of different oxygen depletion indicate that the dominance of Ammonia parkinsoniana over Elphidium spp. (A–E index) was much more pronounced in oxygen-depleted compared to welloxygenated waters. The A–E index has also proven to be a strong, consistent oxygen-stress signal in other coastal areas, for example, Chesapeake Bay and Long Island Sound. The A–E index from sediment cores increased significantly after the 1950s, suggesting increased oxygen stress (in intensity or duration) in the last half century. Buliminella morgani, a hypoxia-tolerant species, known only from the Gulf of Mexico, dominates the present-day population (450%) within areas of chronic seasonal hypoxia, and has also increased markedly in recent decades. Quinqueloculina sp., a hypoxia-intolerant foraminiferan, was a conspicuous member of the fauna from 1700 to 1900, indicating that oxygen
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stress was not a problem prior to 1900, but this species is no longer present on northern Gulf of Mexico shelf in the Louisiana Bight. Multiple lines of evidence from sediment cores indicate an overall increase in phytoplankton productivity and continental shelf oxygen stress (in intensity or duration) in the northern Gulf of Mexico adjacent to the plume of the Mississippi River, especially in the last half of the twentieth century. The changes in these indicators are consistent with the increases in river nitrate-N loading during that same period. OMZ intensity and distribution vary over geological timescales as a result of shifts in productivity or circulation over a few thousands to 10 ky. These changes affect expansions and contractions of the oxygen-depleted waters both vertically and horizontally. Paleoindicators, including foraminiferans, organic carbon preservation, carbonate dissolution, nitrogen isotopes, and Cd:Ca ratios that reflect productivity maxima and shallow winter mixing of the water column, are used to trace longer-term changes in OMZs, similar to studies of continental shelf sediment indicators.
Consequences Direct Effects
The obvious effects of hypoxia/anoxia are displacement of pelagic organisms and selective loss of demersal and benthic organisms. These impacts may be aperiodic so that recovery occurs; may occur on a seasonal basis with differing rates of recovery; or may be permanent so that a shift occurs in long-term ecosystem structure and function. As the oxygen concentration falls from saturated or optimal levels toward depletion, a variety of behavioral and physiological impairments affect the animals that reside in the water column or in the sediments or that are attached to hard substrates (Figure 4). Hypoxia acts as an endocrine disruptor with adverse effects on reproductive performance of fishes, and loss of secondary production may therefore be a widespread environmental consequence of hypoxia. Mobile animals, such as shrimp, fish, and some crabs, flee waters where the oxygen concentration falls below 3–2 mg l 1. As dissolved oxygen concentrations continue to fall, less mobile organisms become stressed and move up out of the sediments, attempt to leave the seabed, and often die (Figure 5). As oxygen levels fall from 0.5 toward 0 mg l 1, there is a fairly linear decrease in benthic infaunal diversity, abundance, and biomass. Losses of entire higher taxa are features of the
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Anoxic sediment, H2S in sediment and water Figure 4 Progressive changes in fish and invertebrate fauna as oxygen concentration decreases from 2 mg l 1 to anoxia. From Rabalais NN, Harper DE, Jr., and Tuner RE (2001) Responses of nekton and demersal and benthic fauna to decreasing oxygen concentrations. In: Rabalais NN and Turner RE (eds.) Coastal and Estuarine Studies 58: Coastal Hypoxia: Consequences for Living Resources and Ecosystems, pp. 115–128. Washington, DC: American Geophysical Union.
depauperate benthic fauna in the severely stressed seasonal hypoxic/anoxic zone of the Louisiana inner shelf in the northern Gulf of Mexico. Larger, longerlived burrowing infauna are replaced by short-lived, smaller surface deposit-feeding polychaetes, and certain typical marine invertebrates are absent from the fauna, for example, pericaridean crustaceans, bivalves, gastropods, and ophiuroids. Long-term trends for the Skagerrak coast of western Sweden in semi-enclosed fiordic areas experiencing increased oxygen stress showed declines in the total abundance and biomass of macroinfauna, abundance and biomass of mollusks, and abundance of suspension feeders and carnivores. These changes in benthic communities result in an impoverished diet for bottom-feeding fish and crustaceans and contribute, along with low dissolved oxygen, to altered sediment biogeochemical cycles. In waters of Scandinavia and the Baltic, there was a reduction of 3 million t in benthic macrofaunal biomass during the worst years of hypoxia occurrence. This loss, however, may be
partly compensated by the biomass increase that occurred in well-flushed organically enriched coastal areas not subject to hypoxia. Where oxygen minimum zones impinge on continental margins or sea mounts, they have considerable effects on benthic assemblages. The benthic fauna of OMZs consist mainly of smaller-sized protozoan and meiofaunal organisms, with few or no macrofauna or megafauna. The few eukaryotic organisms are nematodes and foraminiferans. Meiofauna appear to be more broadly tolerant of oxygen depletion than are macrofauna. The numbers of metazoan meiofaunal organisms, primarily nematodes, are not reduced in OMZs, presumably due to abundant particulate food and reduced predation pressure. In hypoxic waters of the northern Gulf of Mexico, harpacticoid copepod meiofauna are reduced at low oxygen levels, but the nematodes maintain their densities. Benthic macrofauna are found in all hypoxic sediments of the northern Gulf of Mexico, although the density is severely reduced
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Direct mortality
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Figure 5 Effects of hypoxia on fishery resources and the benthic communities that support them. Upper right: Dead demersal and bottom-dwelling fishes killed by the encroachment of near-anoxic waters onto a Grand Isle, Louisiana, beach in August 1990. Photo provided by K. M. St. Pe´. Lower right: dead spider crab (family Majidae) at sediment surface. Photo provided by Franklin Viola. Lower left: dead polychaete (family Spionidae) and filamentous sulfur bacteria. Photo provided by Franklin Viola.
below 0.5 mg l 1, and the few remaining organisms are polychaetes of the families Ampharetidae and Magelonidae and some sipunculans. While permanent deep-water hypoxia that impinges on 2.3% of the ocean’s continental margin may be inhospitable to most commercially valuable marine resources, they support the largest, most continuous reducing ecosystems in the world oceans. Large filamentous sulfur bacteria, Thioploca and Beggiatoa, thrive in hypoxic conditions of 0.1 ml l 1. OMZ sediments characteristically support large bacteria, both filamentous sulfur bacteria and giant spherical sulfur bacteria with diameters of 100–300 mm. The filamentous sulfur bacteria are also characteristic of severely oxygen depleted waters in the northern Gulf of Mexico. Secondary Production
An increase in nutrient availability results in an increase of fisheries yield to a maximal point; then there are declines in various compartments of the fishery as further increases in nutrients lead to seasonal hypoxia and permanent anoxia in semienclosed seas. Documenting loss of fisheries related
to the secondary effects of eutrophication, such as the loss of seabed vegetation and extensive bottom water oxygen depletion, is complicated by poor fisheries data, inadequate economic indicators, increase in overharvesting that occurred at the time that habitat degradation progressed, natural variability of fish populations, shifts in harvestable populations, and climatic variability. Eutrophication often leads to the loss of habitat (rooted vegetation or macroalgae) or low dissolved oxygen, both of which may lead to loss of fisheries production. In the deepest bottoms of the Baltic proper, animals have long been scarce or absent because of low oxygen availability. This area was 20 000 km2 until the 1940s. Since then, about a third of the Baltic bottom area has intermittent oxygen depletion. Lowered oxygen concentrations and increased sedimentation have changed the benthic fauna in the deeper parts of the Baltic, resulting in an impoverished diet for bottom fish. Above the halocline in areas not influenced by local pollution, benthic biomass has increased due mostly to an increase in mollusks. On the other hand, many reports document instances where local pollution resulting in severely depressed oxygen levels has greatly
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impoverished or even annihilated the soft-bottom macrofauna. Eutrophication of surface waters accompanied by oxygen-deficient bottom waters can lead to a shift in dominance from demersal fishes to pelagic fishes. In the Baltic Sea and Kattegatt where eutrophicationrelated ecological changes occurred mainly after World War II, changes in fish stocks have been both positive (due to increased food supply; e.g., pike perch in Baltic archipelagos) and negative (e.g., oxygen deficiency reducing Baltic cod recruitment and eventual harvest). Similar shifts are inferred with limited data on the Mississippi River-influenced shelf with the increase in two pelagic species in bycatch from shrimp trawls and a decrease in some demersal species. Commercial fisheries in the Black Sea declined as eutrophication led to the loss of macroalgal habitat and oxygen deficiency, amid the possibility of overfishing. After the mid-1970s, benthic fish populations (e.g., turbot) collapsed, and pelagic fish populations (small pelagic fish, such as anchovy and sprat) started to increase. The commercial fisheries diversity declined from about 25 fished species to about five in 20 years (1960s to 1980s), while anchovy stocks and fisheries increased rapidly. The point on the continuum of increasing nutrients versus fishery yields remains vague as to where benefits are subsumed by environmental problems that lead to decreased landings or reduced quality of production and biomass.
Future Expectations The continued and accelerated export of nitrogen and phosphorus to the world’s coastal ocean is the trajectory to be expected unless societal intervention takes place (in the form of controls or changes in culture). The largest increases are predicted for southern and eastern Asia, associated with predicted large increases in population, increased fertilizer use to grow food to meet the dietary demands of that population, and increased industrialization. The implications for coastal eutrophication and subsequent ecosystem changes such as worsening conditions of oxygen depletion are significant.
Ecosystems. Fishery Management, Dimension. Upwelling Ecosystems.
Human
Further Reading Dı´az RJ, Nestlerode J, and Dı´az ML (2004) A global perspective on the effects of eutrophication and hypoxia on aquatic biota. In Rupp GL and White MD (eds.) Proceedings of the 7th International Symposium on Fish Physiology, Toxicology and Water Quality, EPA 600/ R-04/049, pp. 1–33. Tallinn, Estonia, 12–15 May 2003. Athens, GA: Ecosystems Research Division, US EPA. Dı´az RJ and Rosenberg R (1995) Marine benthic hypoxia: A review of its ecological effects and the behavioural responses of benthic macrofauna. Oceanography and Marine Biology Annual Review 33: 245--303. Gray JS, Wu RS, and Or YY (2002) Review. Effects of hypoxia and organic enrichment on the coastal marine environment. Marine Ecology Progress Series 238: 249--279. Hagy JD, Boynton WR, and Keefe CW (2004) Hypoxia in Chesapeake Bay, 1950–2001: Long-term change in relation to nutrient loading and river flow. Estuaries 27: 634--658. Helly J and Levin LA (2004) Global distributions of naturally occurring marine hypoxia on continental margins. Deep-Sea Research 51: 1159--1168. Mee LD, Friedrich JJ, and Gomoiu MT (2005) Restoring the Black Sea in times of uncertainty. Oceanography 18: 100--111. Rabalais NN and Turner RE (eds.) (2001) Coastal and Estuarine Studies 58: Coastal Hypoxia – Consequences for Living Resources and Ecosystems. Washington, DC: American Geophysical Union. Rabalais NN, Turner RE, and Scavia D (2002) Beyond science into policy: Gulf of Mexico hypoxia and the Mississippi River. BioScience 52: 129--142. Rabalais NN, Turner RE, Sen Gupta BK, Boesch DF, Chapman P, and Murrell MC (2007) Characterization and long-term trends of hypoxia in the northern Gulf of Mexico: Does the science support the Action Plan? Estuaries and Coasts 30(supplement 5): 753--772. Turner RE, Rabalais NN, and Justic´ D (2006) Predicting summer hypoxia in the northern Gulf of Mexico: Riverine N, P and Si loading. Marine Pollution Bulletin 52: 139--148. Tyson RV and Pearson TH (eds.) Geological Society Special Publication No. 58: Modern and Ancient Continental Shelf Anoxia, 470pp. London: The Geological Society.
See also
Relevant Websites
Coastal Topography, Human Impact on. Ecosystem Effects of Fishing. Eutrophication. Fiordic
http://www.gulfhypoxia.net – Hypoxia in the Northern Gulf of Mexico.
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ICEBERGS D. Diemand, Coriolis, Shoreham, VT, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1255–1264, & 2001, Elsevier Ltd.
Introduction Icebergs are large blocks of freshwater ice that break away from marine glaciers and floating ice shelves of glacial origin. Although they originate on land, they are often included in discussions of sea ice because they are commonly found surrounded by it. However, unlike sea ice, they are composed of fresh water; therefore their origin, crystal structure, and chemical composition, as well as the hazards they pose, are different. They are found in both polar regions, their sizes and numbers generally being greater at higher latitudes. They pose a hazard both to shipping and to seabed structures.
Origins and Spatial Distribution The great ice sheets of Greenland and Antarctica, which produce by far the greatest number of the world’s icebergs, flow off the land and into the sea through numerous outlet glaciers. In many cases, especially in Antarctica, the ice spreads out on the sea surface, staying connected to land and forming a floating ice shelf of greater or lesser extent. There are two major differences between the calving fronts in Greenland and those in Antarctica. First, most of the Greenland icebergs are calved directly from the parent glaciers into the sea, while Antarctic icebergs are mostly calved from the edges of the huge ice shelves that fringe much of the continent. The result is that southern icebergs at the time of calving tend to be very large and tabular, while the northern ones are not so large and have a more compact configuration. Second, the equilibrium line of the Greenland ice sheet is above 1000 m. Therefore, the entire volume of a Greenland iceberg is composed of ice. In Antarctica, on the other hand, the equilibrium line is at or near the edge of the ice shelves, so that icebergs are commonly calved with an upper layer of permeable firn (see Ice Properties below) of varying thickness that influences later deterioration rates and complicates estimates of draft and mass.
The drift of icebergs is largely governed by ocean currents, although wind may exert some influence. Since ocean currents at depth may differ in speed and direction from surface currents, a large iceberg may move in a direction different from that of the surrounding sea ice, creating a patch of open water behind it. Since its speed and direction are heavily dependent on the depth and shape of the keel, which is usually unknown, trajectory predictions are seldom reliable, even when local current profiles are known. Because the Antarctic continent is surrounded by oceans while the Arctic is an ocean surrounded by continents, the drift patterns of the icebergs from these areas are very different. Baffin Bay to North Atlantic Region
About 95% of icebergs in northern latitudes originate on Greenland. Most of these are from western Greenland where they calve directly into Baffin Bay, but a few are produced in eastern Greenland. Many of these remain trapped in the fiords where they originated, deteriorating to a great degree before they reach the sea. Those from eastern Greenland that do reach the sea drift south in the East Greenland Current, a small number continuing south into the North Atlantic where they rapidly dwindle, others being carried around the southern tip of Greenland and then north in the warm West Greenland Current, where the few that survive the long trip join the great numbers of bergs originating from Disko Bay north. Figure 1 shows some of the most active glaciers on Greenland and the drift paths generally followed by icebergs. Icebergs may remain in Baffin Bay for several years, circulating north along the Greenland coast and then south along the Canadian arctic islands. Since the water temperature in Baffin Bay remains consistently low throughout the year, little deterioration takes place. However, many do escape southward through the Davis Strait and drift down the Labrador coast in the cold Labrador Current until they break free of the annual pack ice and reach the Grand Banks off Newfoundland. Icebergs have been sighted as far south as Bermuda and the Azores. Arctic Ocean
The remaining 5% of northern icebergs are calved from numerous glaciers on Ellesmere Island in the
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Baffin Bay
cir cle cti c Ar
60
Di sk o
Ba ffin Is.
˚N
Ba y
Is. Devon
Sc Su ores nd by
p tr u ns e e St
sh av n
Gade
Greenland
Ja ko b
boldt Hum
Up er Ri nav nk ik s
80˚N
Ellesmere Is.
70
˚N
80
˚N
182
Davis strait
70˚N
Labrador Sea Arctic circle
Labrador 60˚N
Figure 1 Sources and drift paths of North Atlantic icebergs.
Canadian Arctic, the many islands in or bordering the Barents, Kara, and Laptev Seas, and Alaska (see Figure 2). Many of these, especially those calved from ice shelves on Ellesmere Island, are tabular in form. When they were first discovered drifting among the Arctic pack, they were referred to as ‘ice islands’, and the name stays with them. Once they have become incorporated into the pack, they tend to stay there indefinitely, although occasionally one may escape and join the southbound flux through Davis Strait or the east coast of Greenland. The sources and trajectories of these icebergs and ice islands are shown in Figure 2. Southern Regions
Since there is no significant runoff from Antarctica, iceberg production accounts for most mass loss from
the continent. Most of these icebergs are calved from the massive ice shelves, such as the Ross, Filchner, Ronne, Larsen, and Amery. About 60–80% by volume are calved from the ice shelves, the remainder from outlet glaciers that empty directly into the sea or from active ice tongues. Once free of the ice front, they drift with the prevailing current along the coast, in some places westward, in others eastward as shown in Figure 3. They may remain close to the coast, where their concentration is the greatest, for periods up to 4 years, protected by the sea ice and the cold water. There are several localized places around the coast where icebergs turn north away from the continent. Once they drift beyond the northern limit of the pack ice, about 601S, they are carried east and north into ever warmer waters until they deteriorate. The most northerly reported sighting was at 261S near the Tropic of Capricorn. Few pass 551S.
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Figure 2 Sources and drift paths of icebergs in the Arctic Ocean.
Numbers and Size Distribution Our knowledge of the numbers and size distribution of icebergs is based on visual observations from ships and aircraft; from radar data from ships, shore, and aircraft; and from satellite imagery. Each method has its advantages and shortcomings. For example, satellite imagery covers very large areas and all times of the year, but will not detect small bergs; ship’s radar will pick up most icebergs within its range but may miss rounded bergs or small bergs in heavy seas; visual observation will catch all sizes of bergs, but
only within a limited area in good weather when someone is looking. Thus, any iceberg census will be slanted toward the size and shape categories favored by the observation method used. The most detailed records of iceberg numbers and sizes in a single location have been kept by the US Coast Guard’s International Ice Patrol (IIP) which was formed in the aftermath of the sinking of the Titanic. The IIP began patrolling the Grand Banks in 1914 and reporting iceberg locations to ships in the area. Since that time the IIP has kept a detailed record of all icebergs crossing 481N. These numbers
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0°E / W 60°S
Antarctic circ le Larsen Ice Shelf
90°W 70°S
Filchner Ice Shelf Ronne Ice Shelf South Pole
Amery Ice Shelf 60°S 90°E
Ross Ice Shelf
70°S 180°E/W
Figure 3 Sources and drift paths of Antarctic icebergs. The shaded box shows the area covered by the satellite image in Figure 5
are highly variable from year to year as is apparent from Figure 4. The reason for this variability is not clear. Both the numbers and sizes are greater at higher latitudes, since the bergs gradually disintegrate as they drift south into warmer waters. No such long-standing record exists for the southern oceans, so estimates of numbers here may be less reliable. However, the National Ice Center, using satellite imagery, does identify and track icebergs whose longest dimension is greater than 10 nautical miles (18.5 km) when first sighted. They also continue to track fragments smaller than this that may break away but are still detectable by satellite radar. At the same time, Norway’s Norsk Polarinstitutt has kept a record of all icebergs sighted in Antarctic waters by ‘ships of opportunity’, which is most ships in the area, since 1981. This data set
includes icebergs of all sizes, but the coverage is restricted to those times and areas where ships are present. Northern Regions
The total estimated volume of ice calved annually from Greenland is about 225765 km3. Estimated numbers of icebergs calved from Greenland’s glaciers range from about 10 000 to 30 000 per year. The greatest numbers in northern oceans are found in Baffin Bay. Icebergs may also be seasonally very numerous along the coast of eastern Canada, especially before the pack ice melts. Of those that drift south into the North Atlantic, the annual numbers crossing 481N, just north of the Grand Banks of Newfoundland, according to the IIP are shown in Figure 4.
Figure 4 Total numbers of icebergs crossing 481N each year from 1900 through 1999. Note: The figures for the years of World War I and II are incomplete. (Data courtesy of the US Coast Guard International Ice Patrol.)
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Once they encounter the warm Gulf Stream waters, they rapidly deteriorate. Southern Regions
The total estimated volume of ice calved annually from Antarctica ranges from 750 to 3000 km3 per year. The occasional release of extremely large icebergs has a major impact on annual estimates of Antarctic mass loss. The iceberg shown in Figure 5. represents as much ice as the total annual mass loss from a ‘normal’ year. The shaded box shown in Figure 3 is the area covered by this image. Estimated numbers of icebergs calved range from 5000 to 10 000 each year.
Shapes and Sizes The range of sizes of icebergs is enormous, spanning about eight orders of magnitude, from small
fragments with a mass around 1000 tonnes to the immense Antarctic tabular begs with masses in excess of 1010 t. Table 1 shows the normal range of iceberg sizes in the Labrador Sea. In terms of shape, no two icebergs are the same. However, as bergs deteriorate they do tend to assume characteristic forms (Figures 6–10). The shape classification in common use is given in Table 2. Specialized terms used in classification and description are defined in the Glossary. It should be borne in mind that the shape or extent of the ‘sail’ does not necessarily reflect the shape of the entire iceberg. Figure 11A and B show a photograph of an iceberg and a computer-generated image of its underwater configuration. The nearly spherical shape of this medium-sized berg suggests that it had rolled, probably recently and probably several times. Any horizontal tongues of ice, or ‘rams’, would have broken away during this energetic process. Figure 12A and B show a larger iceberg that had only tilted from its original in the water. Extensive rams are visible extending outward underwater far beyond the extent of the sail, remnants of a far greater mass that has been lost since the berg moved into relatively warm waters. While these two icebergs have roughly similar shapes above water, their underwater configurations are very different.
Table 1
Figure 5 Satellite image showing an extremely large iceberg breaking away from the Ronne Ice Shelf in 1998. The area covered by this image is indicated in Figure 3. A and B are the remnants of two other very large icebergs, both of which broke free in 1986 and were grounded at the time of this image. C is a rapidly moving stream of ice moving off the continent through the Filchner Ice Shelf and past Berkner Island (D) into the Weddell Sea. (Radarsat data ^ 1998 Canadian Space Agency/Agence spatiale canadienne. Received by the Canada Centre for Remote Sensing (CCRS). Processed by Radarsat International (RSI) and the Alaska SAR facility (ASF). Image enhancement and interpretation by CCRS. Provided courtesy of RSI, CCRS, ASF, and the National Ice Center.)
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Iceberg size categories
Designation
Height (m)
Length (m)
Approximate mass (Mt)
Growler Bergy Bit Small Medium Large Very Large
o1 1–5 5–15 16–45 46–75 475
o5 5–15 15–60 61–120 121–200 4200
0.001 0.01 0.1 2 10 410
Figure 6 Tabular iceberg. (Photograph by Deborah Diemand.)
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Figure 7 Wedge iceberg. (Photograph by Deborah Diemand.) Figure 10 Domed iceberg. (Photograph by Deborah Diemand.)
Table 2
Iceberg shape categories
Designation
Description
Tabular types Tabular Blocky
Nontabular types Wedge
Figure 8 Pinnacle iceberg. (Photograph by Deborah Diemand.) Pinnacle Drydock Dome
Steep sides with a flat top; length to height ratio >5:1 (Figure 6) Similar to tabular, but length to height ratio o5:1
One flat side sloping gradually to the water; the opposite side sloping steeply, the two meeting at the peak as a spine (Figure 7) With one or more sharp peaks (Figure 8) With two or more peaks separated by a water-filled channel (Figure 9) Small with rounded top (Figure 10)
To approximate the mass in tonnes, that of eqn [2] can be used. Mass ¼ 3:01 ½ðlongest sail dimensionðmÞÞ ðorthogonal widthðmÞÞ ðmaximum heightðmÞÞ Figure 9 Drydock iceberg. (Photograph by Deborah Diemand.)
Because of this uncertainty in the underwater shape, it is impossible to calculate iceberg draft and mass accurately using the sail dimensions. However, certain rules of thumb have emerged from empirical studies. To approximate the draft in meters, the relationship of eqn [1] can be used. Draft ¼ 49:4 height0:2
½1
½2
These calculations apply only to icebergs composed entirely of ice, with no firn layer. Northern Regions
In general, the mean size of icebergs in Baffin Bay is about 60 m height, 100 m width and 100 m draft. Mean mass is about 5–10 Mt. The sizes of icebergs in this area are constrained by the water depth near the calving fronts, which is less than 200 m. Icebergs with a mass greater than 20 Mt are extremely rare, and for those found south of 601N, a mass greater
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Figure 11 (A) Medium-sized iceberg grounded in 107 m of water. (B) Computer-generated image showing the underwater profile of this berg. The shape and size of the sail were established through stereophotography; those of the keel were determined using acoustic profiling techniques. These two datasets were joined electronically to create this image as well as Figure 12B. WL, water line. ((A) Photograph by Deborah Diemand; (B) courtesy of Dr. James H. Lever.)
than 10 Mt is seldom found. The maximum sail height on record for an iceberg in the North Atlantic is 168 m. The ice islands in the Arctic Ocean extend about 5 m above sea level. They have a thickness of 30–50 m and an area from a few thousand square metres to 500 km2 or more.
Southern Regions
In general, the thickness of the ice shelves at the calving fronts is about 200–250 m, increasing away from the front. Since the ice edge is very long, and often seaward of seabed obstructions, Antarctic icebergs may be extremely large. The iceberg shown in Figure 5 is more than 5800 km2, larger than the US state of Rhode Island. The
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Figure 12 (A) Large iceberg grounded in 134 m of water. (B) Computer-generated image showing the underwater profile of this berg. WL, water line. ((A) Photograph by Deborah Diemand; (B) courtesy of Dr. James H. Lever.)
largest iceberg ever reported was about 180 km long, with an estimated volume of 1000 km3.
Deterioration Deterioration begins as soon as an iceberg calves from its parent glacier. Even icebergs locked into sea ice over the winter show signs of mass loss. However, significant deterioration does not usually begin until after the berg breaks free from the pack ice and is exposed to warmer surface water and wave action. The major causes of ice loss are melting, calving, and splitting and ram loss. Melting
Ice loss due to melting alone is hard to quantify, but is thought to fall within the shaded region shown in Figure 13. It is highly dependent on water temperature, but is also influenced by wave action, water currents, and bubble release. Melting rate at the
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Antarctic tabular bergs, and the resulting fragments can still be extremely large. In this case, probably the main cause of breakup is flexure due to ocean waves, 160 although grounding or collision, or both, may con140 tribute. It is likely that grounding is the main cause of 120 splitting for bergs smaller than 1 km2. 100 While undercut cliffs on the sides of icebergs tend 80 to shed many small fragments, the corresponding 60 underwater rams remain intact until they are suf40 ficiently large that buoyancy forces alone cause them 20 to break away, or the berg grounds. These rounded 0 fragments, which may be of considerable size, 2 4 8 10 0 12 6 probably represent a large proportion of the domed Seawater temperature (˚C above freezing point) icebergs common in warmer waters. For example, Figure 13 Dependence of the melt rate of icebergs on the the calving of the ram extending to the right in temperature above the freezing point of sea water. Figure 12B would create a new iceberg weighing roughly 100 000 tonnes. Splitting and ram loss are the major cause of size water line is far greater than that over the rest of the ice surface. This causes a groove to form, undercut- reduction in extremely large bergs. ting the ice cliffs and creating sometimes extensive underwater rams. The importance of melting in the overall mass loss Ice Properties depends on the surface/volume ratio, being more Glacial ice is formed by the gradual accumulation of significant for small bergs and growlers than for large snow over many centuries. As the snow compacts ones. and recrystallizes, it forms firn, a granular, permeable One of the side effects of the rapid side melting of material. The firn layer may reach as deep as 100 m icebergs is the vertical mixing of the surrounding in very cold places, but is seldom deeper than 50 m. seawater. Driven mostly by the release of air from the This firn layer is not present on the calving fronts of bubbles in the ice and partly by the lower density of Greenland, but is present on Antarctica’s ice shelves, the fresh water of the melted ice, water flows upward and in the icebergs calved from them. When the firn near the berg, drawing deep water to the surface. The reaches a density of 830 kg m3 the pores close off, combination of nutrients brought up from depth by trapping any air that remains. At this point the ice this process and the decreased salinity of the melt- contains about 10% air by volume. Further densifiwater surrounding the berg results in a specialized cation is a result of compression of the air in the community of plankton and fish in the vicinity of bubbles. The bubbles become smaller and may beicebergs. come incorporated into the crystals through recrystallization. The pressure inside them may be as Calving of Cliff Faces high as 2 MPa (20 bars). 200
_
Melt rate (m y 1)
180
Small pieces of ice are constantly breaking off the sides of icebergs, mostly owing to waterline undercutting. Such calving events may produce only a few small pieces, or a great number, especially in warm water. Usually the individual pieces are quite small and are quickly melted, but the total mass loss can be considerable and the resulting imbalance can cause the berg to roll, causing further ice loss. Once the berg has rolled and stabilized, waterline erosion begins anew. This is probably the major cause of mass loss in medium-sized bergs.
Acoustics
Melting icebergs in the open ocean make a characteristic sound sometimes referred to as ‘bergy seltzer’. This is probably created by the explosion or implosion of bubbles as the ice melts. The frequency range of audible sound produced is quite wide, and is largely masked by ambient ocean noise at frequencies below 6 kHz. The sound seems to vary from berg to berg and is undoubtedly influenced by environmental conditions. Estimated detection distances at frequencies above 6 kHz range from 2 to 150 km.
Splitting and Ram Loss
Splitting occurs when a large iceberg breaks into two or more pieces, each of which is an iceberg in its own right. This is a common occurrence for very large
Ice Temperature
Since ice is a good insulator, the original temperature of a large berg at the time of calving will be retained
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in its central core, and may be as low as 221C. After a year or more in cold water, where little or no ablation takes place, the surface ice will warm to about 01C. In relatively warm waters, however, these outer layers of ice are removed more rapidly than the inner cold core warms up, leaving much colder ice near the iceberg’s surface. Since the strength of ice is greater at lower temperatures, the result of a collision with such a berg could be more severe than with one that had not undergone significant melting. Color
In small quantities, ice appears both colorless and transparent. However, because ice selectively transmits light in the blue portion of the visible spectrum while it absorbs light of other frequencies, sufficiently large pieces of clear, bubble-free ice can appear blue. However, this color is frequently masked in glacial ice by the scattering of light of all wavelengths by the bubbles included in the ice, causing the ice to appear white. Blue bands, commonly present within the greater white mass, are caused when cracks form on the parent glacier or later on the iceberg itself and are filled with meltwater that then freezes relatively bubble free. These blue bands range in size from hairline cracks to a meter or more in width. Green icebergs are fairly common in certain regions. This has variously been attributed to copper or iron compounds, the incorporation of dissolved organic compounds, or to an optical trick caused by red light of the sun near the horizon causing the apparent green color. It is likely that there is no single cause and that all of these factors may make the ice appear green. In some cases the trace substances may originate on land, as in the blue cracks mentioned above. In others they may result from sea water freezing to the underside of an ice shelf. Unlike in ice formed at the water surface, most salts and bubbles are rejected, but certain compounds may be trapped in trace amounts, causing the green appearance of otherwise clear ice. Icebergs may also have bands of brown or black; these are caused by morainal or volcanic material deposited while the ice was still part of the parent glacier.
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much smaller bergs. These ‘small’ bergs may weigh in excess of 100 000 t, but owing to their small above-water size and frequently rounded shape they may not be detected by radar until they are dangerously close to the ship, especially in storm conditions when the radar return from the rough sea surface (sea clutter) will tend to mask the weak radar return from the iceberg. In such conditions the peril is further increased because these relatively small ice masses, tossed by the heavy seas, may reach maximum instantaneous velocities 4–5 times larger than hourly drift speeds. A 4000 t bergy bit moving at the maximum fluid particle velocity of 4.5 m s1 in typical Grand Banks storm waves could have about a third of the kinetic energy of a 1 Mt iceberg drifting at 0.5 m s1 (B1 knot). While the influence of waves on ice movement decreases for larger bergs, icebergs as massive as 1 Mt may still exhibit significantly higher maximum instantaneous velocities than their hourly drift values. Seabed Damage
While small icebergs pose a serious threat to structures at the sea surface, seabed structures such as well-heads, pipelines, cables, and mooring systems are endangered by large icebergs, which may possess a deep enough draft to collide with the seafloor. Marine navigators have long known that the keels of icebergs drifting south over the relatively shallow banks of Canada’s eastern continental shelf may touch the seabed and become grounded. Modern iceberg scours appear in the form of linear to curvilinear scour marks and as pits, and occur from the Baffin Bay/Davis Strait region to the Grand Banks of Newfoundland. They are present at water depths up to about 200 m. Seabed scouring has also been documented in Antarctica, but to date has generated little interest because of the absence of seabed structures. A single scour may be as wide as 30 m, as deep as 10 m, and longer than 100 km. An iceberg may also produce pitting when its draft is suddenly increased through splitting or rolling. It may then remain anchored to the seafloor, rocking and twisting, and may produce a pit deeper than the maximum scour depth. Usage of Icebergs
Economic Importance Hazard to Shipping
A large iceberg poses little threat to shipping on the whole. It will not normally exceed a speed of 1 m s1 (2 knots) and it can be detected with normal marine radar at a considerable distance, allowing the ship to alter course. Ironically, a greater danger is posed by
In the past there has been considerable interest in the possibility of transporting icebergs, representing as they do an essentially unlimited supply of fresh water, to arid areas such as Saudi Arabia, Western Australia, and South America. The two seemingly insurmountable problems that need to be solved are propulsion and prevention of in-transit breakup in
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warm seas. Proposed means of moving a sufficiently large ice mass over such long distances have ranged from conventional towing to use of a nuclear submarine to wind power. None has proven feasible. Destruction of Icebergs
Attempts at destroying icebergs have been numerous and varied. Perhaps the most-studied technique has involved the use of explosives, which have been extensively tested on glacial ice in the form of glaciers and ice islands. Both crater blasting and bench blasting have been attempted. The results of this testing suggest that ice is as difficult to blast as typical hard rock, and that therefore the use of explosives for its destruction is impractical. Other methods tested include spreading carbon black on the berg’s surface to accelerate melting, and introducing various gases into the ice to create holes or cavities that can then be filled with explosives of choice. Attempts have also been made to cut through the ice using various means. There is little evidence that any great success was achieved with any of these methods. The only report of a successful attempt to break up an iceberg involved the use of thermit, a welding compound that reacts at very high temperatures (B30001C). The explanation was that the very high heat produced by the thermit caused massive thermal shock within the mass of ice that ultimately resulted in its disintegration, much as glass can be fragmented by extreme temperature changes.
Conclusions There is a great deal of uncertainty surrounding iceberg properties, behavior, drift, and other aspects relating to individual icebergs as opposed to laboratory samples or intact glaciers. This is mostly because of the high cost of expeditions to the remote areas where icebergs are most numerous, and the inherent dangers of hands-on measurement and sampling.
Equilibrium line On a glacier, the line above which there is a net gain due to snow accumulation and below which there is a net loss due to melt. Firn Permeable, partially consolidated snow with density between 400 kg m 3 and 830 kg m 3. Growler A small fragment of glacial ice extending less than a meter above the sea surface and having a horizontal area of about 20 m2. Keel The underwater portion of an iceberg. Ram Lobe of the underwater portion of an iceberg that extends outward, horizontally, beyond the sail. Sail The above-water portion of an iceberg.
See also Antarctic Circumpolar Current. Arctic Ocean Circulation. Current Systems in the Southern Ocean. Florida Current, Gulf Stream, and Labrador Current. Ice-induced Gouging of the Seafloor. Sea Ice: Overview. Sea Ice. Sonar Systems. Sub Ice-Shelf Circulation and Processes. Weddell Sea Circulation. Wind Driven Circulation.
Further Reading Colbeck SC (ed.) (1980) Dynamics of Snow and Ice Masses. New York: Academic Press. Husseiny AA (ed.) (1978) First International Conference on Iceberg Utilization for Fresh Water Production, Weather Modification, and Other Applications. Iowa State University, Ames, 1977. New York: Pergamon Press. Vaughan D (1993) Chasing the rogue icebergs. New Scientist, 9 January. International Ice Patrol (IIP): http://www.uscg.mil/lantarea /iip/home.html Library of Congress Cold Regions bibliography: http:// lcweb.loc.gov/rr/scitech/coldregions/welcome.html National Ice Center (NIC): http://www.natice.noaa.gov/
Glossary Calving The breaking away of an iceberg from its parent glacier or ice shelf. Also the subsequent loss of ice from the iceberg itself.
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ICE-INDUCED GOUGING OF THE SEAFLOOR W. F. Weeks, Portland, OR, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1264–1270, & 2001, Elsevier Ltd.
Introduction Inuit hunters have long known that both sea ice and icebergs could interact with the underlying sea floor, in that sea floor sediments could occasionally be seen attached to these icy objects. As early as 1855, no less a scientific luminary than Charles Darwin speculated that gouging icebergs could traverse isobaths, concluding that an iceberg could be driven over great inequalities of [a seafloor] surface easier than could a glacier. Only in 1924 did similar inferences of the possibility of sea ice and icebergs affecting the seafloor again begin to reappear in the scientific literature. However, direct observations of the nature of these interactions were lacking until the early 1970s. The reason for the increased interest in this rather exotic phenomena was initially the discovery of the supergiant oil field on the edge of the Beaufort Sea at Prudhoe Bay, Alaska. Because the initial successful well was located on the coast, there was immediate interest in the possibility of developing offshore fields to the north of Alaska and Canada. It was also apparent that if major oil resources occurred to the north of Alaska and Canada, similar resources might be found on the world’s largest continental shelf located to the north of the Russian mainland. Offshore wells are typically tied together by subsea pipelines which take the oil from the individual wells either to a central collection point where tanker pickup is possible or to the coast where the oil can be fed into a pipeline transportation system. If sea ice processes could result in major distrubances of the seafloor, this would clearly become a major consideration in pipeline design. Some time after the Prudhoe Bay find, another large offshore oil discovery was made to the east of Newfoundland. Here the ice-induced gouging problem was caused not by sea ice but by icebergs. In both cases the initial step in treating these perceived problems was to investigate the nature of these ice–seafloor interactions. One needs to know the frequency of gouging events in time and space as well as the widths and depths of the gouges, the water depth range in which this phenomenon occurs, and the effective lifetime of a gouge after its initial formation. Also of importance are the type of subsea soil
and the nature and extent of the soil movements below the gouges, as this information is essential in calculating safe burial depths for subsea structures. The following attempts to summarize our current knowledge of this type of naturally occurring phenomenon. As extensive studies of ice-induced disruptions of the seafloor have been undertaken only during the last 30 years, there is as yet no commonly agreed upon terminology for this phenomenon. In the literature the process has been described as scouring, scoring, plowing and gouging. Here gouging is used because at least in the case of sea ice, it is felt that it more accurately describes the process than do the other terms.
Observational Techniques A variety of techniques have been used to study the gouging phenomenon. Typically, a fathometer is used to resolve the seafloor relief directly beneath the ship with a precision of better than 10 cm (Figure 1) while, at the same time, a sidescan sonar system provides a sonar map of the seafloor on either side of the ship (Figure 2). Total sonar swath widths have been typically 200 to 250 m not including a narrow area directly beneath the ship that was not imaged. The simultaneous use of these two different types of records allows one to both measure the depth and width of the gouge (fathometer) at the point where it is crossed by the ship track and also to observe the general orientation and geometry of the gouge track (sonar). Along the Alaskan coast the ships used have been small, allowing them to operate in shallow water. In some cases divers have been used to examine the gouges and, at a few deeper water sites off the Canadian east coast, manned submersibles have been used to gather direct observations of the gouging process.
Results Arctic Shelves
The most common features gouging the shelves of the Arctic Ocean are the keels of pressure ridges that are made of deformed sea ice. As the ice pack moves under the forces exerted on it by the wind and currents, pressure ridges typically form at floe boundaries as the result of differential movements between the floes. These features can be very large. Pressure
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ICE-INDUCED GOUGING OF THE SEAFLOOR
0
30
Depth (m)
35
40
15
30
45 46
1
16
31
2
17
32
47
3
18
33
48
4
19
34
49
5
20
35
50
6
21
36
51
22
37
52
8
23
38
53
9
24
39
54
25
40
55
10
26
41
27
42
57
43 44
45
15
0
100
200
300 Distance (m)
400
500
Figure 1 Fathogram of an ice-gouged seafloor. Water depth is 36 m. Record taken 25 km NE of Cape Halkett, Beaufort Sea, offshore Alaska. The multiple reflections from the upper layer of the seafloor are the result of the presence of a thawed active layer in the subsea permafrost. (From Weeks et al., 1983.)
125 100 75
Distance across track (m)
50 25 0 0 25 50 75 100 125 0
100
200
300
400
500
Distance along ship's track (m) Figure 2 Sonograph of an ice-gouged seafloor. Water depth is 20 m. Record taken 20 km NE of Cape Halkett, Beaufort Sea, offshore Alaska. (From Weeks et al., 1983.)
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ICE-INDUCED GOUGING OF THE SEAFLOOR
ridge keels with drafts up to 50 m have been observed in the upward-looking sonar records collected by submarines passing beneath the ice. The distribution of keel depths of these deformed features is approximately a negative exponential, in that there are many shallow ridge keels while deep keels are rare. Deformation can be particularly intense in nearshore regions where the moving ice pack contacts an immovable coast with the formation of large shear ridges that can extend for many kilometers. Such ridges are frequently anchored at shoal areas where large accumulations of highly deformed sea ice can build up. Such grounded ice features, which are referred to using the Russian term stamuki, can have freeboards of as much as 10 m and lateral extents of 10–15 km. As the offshore ice field moves against and along the coast, it exerts force on the sides of any such grounded features, causing them to scrape and plough their way along the seafloor. Considering that surficial sediments along many areas of the Beaufort Sea Coast of Canada and Alaska are fine-grained silts, it is hardly surprising that over a period of time such a process can cause extensive gouging of the seafloor sediments. Figure 3 is a photograph showing active gouging. The relatively undeformed first-year sea ice in the foreground is moving away from the viewer and pushing against a piece of grounded multiyear sea ice (indicated by its rounded upper surface formed during the previous summer’s melt period), pushing it along the coast. The interaction between the first-year and the multiyear ice has resulted in a pileup of broken first-year ice, which when the photograph was taken was higher than the upper surface of the multiyear ice. Also evident is the track cut in the first-year ice as it moves past the
Figure 3 Photograph of active ice gouging occurring along the coast of the Beaufort Sea. The grounded multiyear ice floe that is being pushed by the first-year ice has a freeboard of B2 m. The thickness of the first-year ice is B0.3 m. (Photograph by GFN Cox.)
193
multiyear ice. The fact that both the multiyear ice and the pileup of first year ice are interacting with the seafloor is indicated by the presence of bottom sediment in the deformed first-year ice and on the far side of the multiyear ice. The maximum water depth in which contemporary sea ice gouging is believed to occur is roughly 50–60 m, corresponding approximately to the draft of the largest pressure ridges. Although gouges occur in water up to 80 m deep, it is generally believed that these deep-water gouges are relicts that formed during periods when sea level was lower than at present. Not surprisingly the depths of gouges in the seafloor mirrors the keel depth distributions observed in pressure ridge keels in that they are also well approximated by a negative exponential with the character of the falloff as well as the number of gouges varying with water depth. As with ridge keels, shallow gouges are common and deep gouges are rare. As might be expected, there are fewer deep gouges in shallow water as the large ice masses required to produce them have already grounded farther out to sea. For instance, along the Beaufort Coast in water 5 m deep, a 1 m gouge has an exceedance probability of approximately 10 4; that is, 1 gouge in 10 000 will on the average be expected to have a depth equal to or greater than 1 m. In water 30 m deep, a 3.4 m gouge has the same probability of occurrence. Gouges in excess of 3 m deep are not rare and 8 m gouges have been reported in the vicinity of the Mackenzie Delta. As might be expected, the nature of gouging varies with changes in seafloor sediment types. Along the Beaufort coast where there are two distinct soil types, the gouges in stiff, sandy, clayey silts are typically more frequent and slightly deeper than those found in more sandy sediments. Presumably the gouges in the more sandy material are more easily obliterated by wave and current action. It is also reasonable to assume that the slightly deeper gouges in the silts provide a better picture of the original incision depths. An extreme example of the effect of bottom sediment type on gouging can be found between Sakhalin Island and the eastern Russian mainland. Here the seafloor is very sandy and the currents are very strong. As a result, even though winter observations have conclusively shown that seafloor gouging is common, summer observations reveal that the gouges have been completely erased by infilling. Gouge tracks in the Beaufort Sea north of the Alaskan and western Canadian coast generally run roughly parallel to the coast, with some regions having an excess of 200 gouges per kilometer.
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ICE-INDUCED GOUGING OF THE SEAFLOOR
Histograms of the frequency of occurrence of distances between gouges are also well described by a negative exponential, a result suggesting that the spatial occurrence of gouges may be described as a Poisson process. Gouge occurrence is hardly uniform, however, with significantly higher concentrations of gouges occurring on the seaward sides of shoal areas and barrier islands and fewer gouges on their more protected lee sides. This is clearly shown in Figure 4 in that the degree of exposure to the moving pack ice decreases in the sequence Jones Islands–Lonely–Harrison Bay–Lagoons. Figure 5 is a sketch based on divers’ observations showing gouging along the Alaskan coast of the Beaufort Sea. The second features producing gouges on the surface of the outer continental shelf of the Arctic Ocean are ice islands. These features are, in fact, an unusual type of tabular iceberg formed by the gradual breakup of the ice shelves located along the north coast of Ellesmere Island, the northernmost of the Canadian Arctic Islands. Once formed, an ice island can circulate in the Arctic Ocean for many years. For
4
10
3
Number of gouges observed
10
2
10
n = 16 620
instance, the ice island T-3 drifted in the Arctic Ocean between 1952 and 1979, ultimately completing three circuits of the Beaufort Gyre (the large clockwise oceanic circulation centered in the offshore Beaufort Sea) before exiting the Arctic Ocean through Fram Strait between Greenland and Svalbard (Spitzbergen). The lateral dimensions of ice islands can vary considerably from a few tens of meters to over ten kilometers. Thicknesses, although variable, are typically in the 40–50 m range; ice islands possess freeboards in the same range as exhibited by larger sea ice pressure ridges. In the study of fathometer and sonar data on gouge distributions and patterns, no attempt is usually made to separate gouges made by pressure ridges from gouges made by ice islands, as the ice features that made the gouges are commonly no longer present. However, it is reasonable to assume that many of the very wide, uniform gouges are the result of ice islands interacting with the seafloor. It is relatively easy to characterize the state of the gouging existing on the seafloor at any given time. It is another matter to answer the question of how deep one must bury a pipeline, a cable, or some other type of fixed structure beneath the seafloor to reduce the chances of it being impacted by moving ice to some acceptable level. To answer this question one needs to know the rates of occurrence of new gouges; these values are in many locations still poorly known, as they require replicate measurements over a period of many years so that new gouges can be counted and the rate of infilling of existing gouges can be estimated. The best available data on gouging occurrence along the Beaufort Coast has been collected by the Canadian and US Geological Surveys and indicates that gouging occurs in rather large scale regional events related to severe storms that drive the
1
10
2869 842 41
0
10
10
_1
0
1
2
3
Gouge depth (m) Figure 4 Semilogarithmic plot of the number of gouges observed versus gouge depth for four regions along the Alaskan coast of the Beaufort Sea. , Jones Island and east; W, Lonely;, n Harrison Bay; J, Lagoons. (From Weeks et al., 1983.)
Figure 5 Diver’s sketch of active gouging occurring off the coast of the Alaskan Beaufort Sea. Some sense of the scale of this drawing can be gained from the observation that commonly the thickness of the ice blocks in such ridges ranges between 0.3 and 1.0 m. (Drawing by TR Alpha, US Geological Survey.)
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ICE-INDUCED GOUGING OF THE SEAFLOOR
pack ice inshore – about once every 4–5 years. Because there is no absolute technique for dating the age of existing gouges, there is no current method for determining whether a particular gouge that one observes on the seafloor formed during the last year or sometime during the last 6000 years (after the Alaskan shelf was submerged as the result of rising sea levels at the end of the Pleistocene). It is currently believed that gouges in shallow water formed quite recently. For instance, during the late summer of 1977 when the Beaufort Coast was comparatively ice free, strong wave action during late summer storms obliterated gouges in water less than 13 m deep and caused pronounced infilling of gouges in somewhat deeper water. It is generally estimated that such storms have an average recurrence interval of approximately 25 years. On the other hand, gouges in water deeper than B60 m are believed to be relatively old in that they are presumed to be below the depth of currently active gouging. The lengths of time represented by the gouges observed in water depths between 20 and 50 m are less well known. Unfortunately, this is the water depth range in which the largest gouges are found. Stochastic models of gouging occurrence, which incorporate approximate simulations of subsea sediment transport, suggest that if seafloor currents are sufficiently strong to exceed the threshhold for sediment movement, gouge infilling will occur comparatively rapidly (within a few years). To date, two different procedures have been used to estimate the depth of gouges with specified recurrence intervals. In the first case, the available rates and geometric characteristics of new gouges at a particular site are used. If a multiyear high-quality dataset of new gouge data is available for the region of interest, this would clearly appear to be the favored approach. In the other case, information on the draft distribution of pressure ridge keels and pack ice drift rates are combined to calculate the rates of gouging. The problem with using pack ice drift rates and pressure ridge keel depths is that drift rates in the near shore where grounding is occurring are undoubtedly less than in the offshore where the ice can move relatively unimpeded. In addition, it is doubtful that offshore keel depths observed in water sufficiently deep to allow submarine operations provide accurate estimates of keel depths in the near-shore. One might suggest that offshore engineers should simply bury pipelines and cables at depths below those of known gouges and forget about all these statistical considerations. Although this sounds attractive, it is not a realistic resolution of the problem. In the first place, for a pipeline to avoid deformation, buckling, and failure it must, depending on the
195
nature of the seafloor sediments, be buried at depths up to two times that of the design gouging event. Considering that gouges in excess of 3 m are not rare at many locations along the Beaufort Coast of Alaska and Canada, this means burial depths 45 m. Burial at such depths is extremely costly and time consuming considering the very short operating season in the offshore Arctic. It is also near the edge of existing subsea trenching technology. Another factor to be considered here is that the deeper the pipeline is buried, the more likely it is that its presence will result in the thawing of subsea permafrost that is known to exist in many areas of the Beaufort and Chukchi Shelves. Such thawing could result in settlement in the vicinity of the pipe, which could threaten the integrity of the pipeline. There are engineering remedies to this problem but, as one might expect, they are expensive. Canadian East Coast
The primary problem to the east of Newfoundland and off the coast of Labrador is not sea ice but icebergs. Although the majority of these originate from the Greenland Ice Sheet calving to the west into Baffin Bay, some icebergs do originate from the East Greenland coast and from the Canadian Arctic (see Icebergs). Considering their great size and deep draft, icebergs are formidable adversaries. Iceberg gouges with lengths 460 km are known and many have lengths 420 km. Gouge depths can exceed 10 m and recent gouging is known to occur at depths of at least 230 m along the Baffin and Labrador Shelves. Icebergs also appear to be able to traverse vertical ranges of bathymetry up to at least 45 m. As with the sea-ice-induced gouges of the Arctic shelves, iceberg gouge depths are exponentially distributed, with small gouges being common and deep gouges being rare. Iceberg gouges can be straight or curved. Pits are also common, which occur when the iceberg draft is suddenly increased through splitting and rolling. The iceberg can then remain fixed to the seafloor while it rocks and twists as a result of the wave and current forces that impinge upon it. At such times pits can be produced that are deeper than the maximum gouge depth otherwise associated with the iceberg. Although considerable information is available on the statistics of iceberg gouge occurrences, in studies of iceberg groundings more attention has been paid to the energetics of the ice–sediment interactions than to the statistics. Also, more emphasis has been placed on iceberg ‘management’ in order either to reduce or to remove the threat to a specific offshore operation. For instance, as icebergs are discrete
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ICE-INDUCED GOUGING OF THE SEAFLOOR
objects as compared to pressure ridges, which are giant piles of ice blocks that are frequently poorly cemented together, icebergs can be deflected away from a production site via the use of tugboats in combination with trajectory modeling. Such schemes can also utilize aerial, satellite, and ship reconnaissance techniques to provide operators with an adequate warning of a potential threat. Surface-based over-the-horizon radar technology has also been developed as an iceberg reconnaissance tool, but at the time of writing it has not been used operationally.
quantitative understanding of processes such as ice gouging becomes essential for safe development. Lack of such understanding leads to poor, inefficient designs and the increased possibility of failures. The last thing that either the environmental community or the petroleum industry needs is an offshore oil spill in the high Arctic.
See also Icebergs. Rigs and offshore Structures. Sea Ice: Overview. Sea Ice. Sub Ice-Shelf Circulation and Processes. Sub-sea Permafrost.
Other Locations
Although the discussion here has primarily focused on the Beaufort Coast of the Arctic Ocean and the offshore regions of Newfoundland and Labrador, this is only because the interest in offshore oil and gas reserves at these locations has resulted in the collection of observational information on the local nature of the ice-induced gouging phenomenon. However, regions where ice-induced gouging is presently active are very large and include the complete continental shelf of the Arctic Ocean, the continental shelves of Greenland, of eastern Canada, and of Svalbard, as well as the continental shelf of the Antarctic continent where typical shelf icebergs are known to have drafts of B200 m with lateral dimensions as large as 100 km. In addition, relict iceberg gouges exist and as a result affect seafloor topography in regions where icebergs are no longer common or no longer occur, such as the Norwegian Shelf and even the northern slope of Little Bahama Bank and the Straits of Florida. Finally, although 230 m is a reasonable estimate for the maximum depth of active iceberg gouging, relict gouges presumed to have occurred during the Pleistocene have been discovered at depths from 450 m to at least 850 m on the Yermak Plateau located to the northwest of Svalbard in the Arctic Ocean proper.
Conclusions There is clearly much more that we need to know concerning processes acting on the poorly explored continental shelves of the polar and subpolar regions. The ice-induced gouging phenomenon is clearly one of the more important of these, in that an understanding of this process is essential to both the engineering and the scientific communities. At first glance, icebergs and stamuki zones might appear to be exotic entities that are out of sight somewhere way to the north or south and that can be safely put out of mind. However, as oil production moves to ever more difficult frontier areas such as the offshore Arctic, a
Further Reading Barnes PW, Schell DM, and Reimnitz E (eds.) (1984) The Alaskan Beaufort Sea: Ecosystems and Environments. Orlando: Academic Press. Colony R and Thorndike AS (1984) An estimate of the mean field of arctic sea ice motion. Journal of Geophysical Research 89(C6): 10623--10629. Darwin CR (1855) On the power of icebergs to make rectilinear, uniformly-directed grooves across a submarine undulatory surface. London, Edinburgh, and Dublin Philosophical Magazine and Journal of Science 10: 96--98. Goodwin CR, Finley JC and Howard LM (1985) Ice Scour Bibliography. Environmental Studies Revolving Funds Rept. No. 010, Ottawa. Lewis CFM (1977) The frequency and magnitude of drift ice groundings from ice-scour tracks in the Canadian Beaufort Sea. Proceedings of 4th International Conference on Port Ocean Engineering Under Arctic Conditions. Newfoundland: Memorial University. vol. 1, 567–576. Palmer AC, Konuk I, Comfort G, and Been K (1990) Ice gouging and the safety of marine pipelines. Offshore Technology Conference, Paper 6371, 235--244. Reed JC and Sater JE (eds.) (1974) The Coast and Shelf of the Beaufort Sea. Arlington, VA: Arctic Institute of North America. Reimnitz E, Barnes PW, and Alpha TR (1973) Bottom features and processes related to drifting ice, US Geological Survey Miscellaneous Field Studies, Map MF532. Washington, DC: US Geological Survey. Vogt PR, Crane K, and Sundvor E (1994) Deep Pleistocene iceberg plowmarks on the Yermak Plareau: sidescan and 3.5 kHz evidence for thick calving ice fronts and a possible marine ice sheet in the Arctic Ocean. Geology 22: 403--406. Wadhams P (1988) Sea ice morphology. In: Leppa¨ranta M (ed.) Physics of Ice-Covered Seas, pp. 231--287. Helsinki: Helsinki University Printing House. Weeks WF, Barnes PW, Rearic DM, and Reimnitz E (1983) Statistical aspects of ice gouging on the Alaskan Shelf of the Beaufort Sea. Cold Regions Research and Engineering
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ICE-INDUCED GOUGING OF THE SEAFLOOR
Laboratory Report. New Hampshire, USA: Hanover. 83–21. Weeks WF, Tucker WB III, and Niedoroda AW (1985) A numerical simulation of ice gouge formation and infilling on the shelf of the Beaufort Sea. Proceedings, International Conference on Port Ocean Engineering
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Under Arctic Conditions, Narssarssuaq, Greenland 1: 393--407. Woodworth-Lynas CMT, Simms A, and Rendell CM (1985) Iceberg grounding and scouring on the Labrador continental shelf. Cold Regions Science and Technology 10(2): 163--186.
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ICE–OCEAN INTERACTION J. H. Morison, University of Washington, Seattle, WA, USA M. G. McPhee, McPhee Research Company, Naches, WA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1271–1281, & 2001, Elsevier Ltd.
Introduction The character of the sea ice cover greatly affects the upper ocean and vice versa. In many ways icecovered seas provide ideal examples of the planetary boundary layer. The under-ice surface may be uniform over large areas relative to the vertical scale of the boundary layer. The absence of surface waves simplifies the boundary layer processes. However, thermodynamic and mechanical characteristics of ice–ocean interaction complicate the picture in unique ways. We discuss a few of those unique characteristics. We deal first with how momentum is transferred to the water and introduce the structure of the boundary layer. This will lead to a discussion of the processes that determine the fluxes of heat and salt. Finally, we discuss some of the unique characteristics imposed on the upper ocean by the larger-scale features of a sea ice cover.
Drag and Characteristic Regions of the Under-ice Boundary Layer To understand the interaction of the ice and water, it is useful to consider three zones of the boundary layer: the molecular sublayer, surface layer, and outer layer (Figure 1). Under a reasonably smooth and uniform ice boundary, these can be described on the basis of the influence of depth on the terms of the equation for a steady, horizontally homogeneous boundary layer (eqn [1]). ifV ¼
q qV q qV n þ K r1 rh p qz qz qz qz
½1
The coordinate system is right-handed with z positive upward and the origin at the ice under-surface. V is the horizontal velocity vector in complex notation (V ¼ u þ iv), r is water density, and p is pressure. An eddy diffusivity representation is used for turbulent
198
shear stress, KðqV=qzÞ ¼ V 0 w0 , where K is the eddy diffusivity. The term nðqV=qzÞ is the viscous shear stress, where n is the kinematic molecular viscosity. The pressure gradient term, r1rhp is equal to r1 ðqp=qx þ iqp=qyÞ. The stress gradient term due to molecular viscosity is of highest inverse order in z. It varies as z2, and therefore dominates the stress balance in the molecular sublayer (Figure 1) where z is vanishingly small. As a result the viscous stress, nðqV=qzÞ, is effectively constant in the molecular sublayer, and the velocity profile is linear. The next layer away from the boundary is the surface layer. Here the relation between stress and velocity depends on the eddy viscosity, which is proportional to the length scale and velocity scale of turbulent eddies. The length scale of the turbulent eddies is proportional to the distance from the boundary, |z|. Therefore, the turbulent stress term varies as z1 and becomes larger than the viscous term beyond z greater than (1/k)(n/u*0 ), typically a fraction of a millimeter. The velocity scale in the surface layer is u*0 , where ru2*0 is equal to t0, the average shear stress at the top of the boundary layer. Thus, K is equal to ku*0 |z|, where Von Ka¨rma¨n’s constant, k, is equal to 0.4. Because the turbulent stress term dominates the equations of motion, the stress is roughly constant with depth in the surface layer. This and the linear z dependence of the eddy coefficient result in the log-layer solution or ‘law of the wall’ (eqn [2]). u 1 1 z ¼ ln z þ C ¼ ln u* k k z0
½2
C ¼ (ln z0)/k is a constant of integration. Under sea ice the surface layer is commonly 1–3 m thick. The surface layer is where the influence of the boundary roughness is imposed on the planetary boundary layer. In the presence of under-ice roughness, the average stress the ice exerts on the ocean, t0, is composed partly of skin friction due to shear and partly of form drag associated with pressure disturbances around pressure ridge keels and other roughness elements. Observations under very rough ice have shown a decrease in turbulent stress toward the surface, presumably because more of the momentum transfer is taken up by pressure forces on the rough surface. The details of this drag partition are not known. Drag partition is complicated further for cases in which stratification exists at depths shallow
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ICE–OCEAN INTERACTION
〈w ′T ′〉0 = T dT / dz 〈w ′S ′〉0 = S dS/dz
N NV ≈ 0 Nz Nz
(
Molecular sublayer
U = du /dz
z0 = hs /30
N K NV ≈ 0 N z Nz
(
Ice
0
Z0
(
S
(
Ice
T
199
23°
= u*20
Depth below ice (m) for u *0 = 0.01 m s
s|
v /u *0
|u| = 13.5u*0 45°
u / u*0
_ 0.1
Vice
10 _ 0.2
15
45° Outer layer
20
ifV ≈ N K NV Nz Nz
(
(
hs
_1
Surface layer
Z s| 5
Stress and velocity vectors in plan view
_ 0.3
25
= zf /u*0 _5
_ 0.4 0
5
10
15
u / u*0 and v /u *0 Figure 1 Illustration of three regions of the planetary boundary layer under sea ice: molecular sublayer, surface layer, and outer layer. The velocity profiles are from the Rossby similarity solution (eqns [8], [9], [10] and [11]) for u*0 ¼ 0.01 m s1, z0 ¼ 0.06 m, Z* ¼ 1. The stress and velocity vector comparisons are from the same solution.
compared to the depth of roughness elements. Then it also becomes possible to transfer momentum by internal wave generation. However, for many purposes t0 is taken as the turbulent stress evaluated at z0. Laboratory studies of turbulent flow over rough surfaces suggest that z0 may be taken equal to hs/30, where hs is the characteristic height of the roughness elements. In rare situations the ice surface may be so smooth that bottom roughness and form drag are not factors in the drag partition. In such a hydrodynamically smooth situation, the turbulence is generated by shear induced instability in the flow. The surface length scale, z0, is determined by the level of turbulent stress and is proportional to the molecular sublayer thickness according to the empirically derived relation z0 ¼ 0.13(n/u*0 ). In the outer layer farthest from the boundary, the Coriolis and pressure gradient terms in eqn [1], which have no explicit z dependence, are comparable to the turbulent stress terms. The presence of the Coriolis term gives rise to a length scale, h, for the outer boundary layer equal to u*0 /f under neutral stratification. This region is far enough from the boundary so that the turbulent length scale becomes
independent of depth and in neutral conditions has been found empirically to be l ¼ xnu*0 /f, where xn is 0.05. For neutral stratification, u* and h are the independent parameters that define the velocity profile over most of the boundary layer. The ratio of the outer length scale to the surface region length scale, z0, is the surface friction Rossby number, R0 ¼ u*0 /(z0f). Solutions for the velocity in the outer layer can be derived for a wide range of conditions if we nondimensionalize the equations with these Rossby similarity parameters, u*0 /f and u*0 . However, the growth and melt of the ice produce buoyancy flux that strongly affects mixing. Melting produces a stabilizing buoyancy flux that inhibits turbulence and contracts the boundary layer. Freezing causes a destabilizing buoyancy flux that enhances turbulence and thickens the boundary layer. We can account for the buoyancy flux effect by adjusting the Rossby parameters dealing with length scale. We define the scale of the outer boundary layer as hm ¼ u*0 Z* =f . If the mixing length of the turbulence in the outer layer is lm ¼ xn u*0 Z2* =f , it interpolates in a reasonable way between known values of lm for neutral stratification (xnu*0 /f ) and stable stratification (RcL) if Z* is given
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200
ICE–OCEAN INTERACTION
as eqn [3]. Z* ¼
x u 1 1=2 1þ n * f Rc L
In the outer layer, eqns [5] and [8] are satisfied for nondimensional velocity given by eqn [10]. ½3
ˆ U ¼ idˆ ed z
for zrz0
½10
Rc is the critical Richardson number; the Obukhov length, L, is the ratio of shear and buoyant production of turbulent energy, ru3*0 =kg/r0 w0 S; and /r0 w0 Sg/r is the turbulent buoyancy flux. With this Rossby similarity normalization of the equations of motion, we can derive analytical expressions for the under-ice boundary layer profile that are applicable to a range of stratification. For large |z|, V will approach the free stream geostrophic velocity, V¯ g ¼ Ug þ iVg ¼ f 1 r1 rh p. Here we will assume this is zero. However, surface stressdriven absolute velocity solutions can be superimposed on any geostrophic current. We also ignore the time variation and viscous terms and define a normalized stress equal to S ¼ ðKqV=qzÞ=u2*0 . The velocity is nondimensionalized by the friction velocity and the boundary layer thickness, U ¼ Vfhm =u2*0 , and depth is nondimensionalized by the boundary layer thickness scale, z ¼ z/hm. With these changes eqn [1] becomes eqn [4].
Thus the velocity is proportional to stress but rotated 451 to the right. As we see in the derivation of the law of the wall [2], the surface layer variation of the eddy viscosity with depth is critical to the strong shear present there. Thus eqn [10] will not give a realistic profile in the surface layer. We define the nondimensional surface layer thickness, zsl, as the depth where the surface layer mixing length, |z|, becomes equal to the outer layer mixing length, lm ¼ xn u*0 Z2* =f . We find zsl is equal to Z* xn and applying the definition [6] gives K* as K*sl ¼ kz/Z* in the surface layer. If we approximate the stress profile [8] by a Taylor series, we can integrate [5] with K*sl substituted for K* to obtain the velocity profile in the surface layer.
iU ¼ qS=qz
Eqn [11] is analogous to [2] except for the introduction of the dˆðzsl zÞ term. This is the direct result of accounting for the stress gradient in the surface layer. This term is small compared to the logarithmic gradient. Figure 1 illustrates the stress and velocity vectors at various points in the boundary layer as modeled by eqns [8] through [11]. For neutral conditions the nondimensional boundary layer thickness is typically 0.4 (dimensional thickness is 0.4u* /f). Through the outer layer, the velocity vector is 451 to the right of the stress vector as a consequence of the idˆ peið451Þ multiplier in [10]. As the ice surface is approached through the surface layer, the stress vector rotates 10–201 to the left to reach the surface direction. However, in the surface layer the velocity shear in the direction of the surface stress is great because of the logarithmic profile. Thus, as the surface is approached, the velocity veers to the left twice as much as stress. At the surface the velocity is about 231 to the right of the surface stress. It is commonly useful to relate the stress on underice surface to the relative velocity between ice and water a neutral-stratification drag coefficient, 2 where Cz is the drag coefficient for ru2*0 ¼ rCz VðzÞ depth z. If z is in the log-layer, eqn [2] can be used to derive the relation between ice roughness and the drag coefficient. We find that Cz ¼ k2[ln(z/z0)]2.
½4
In terms of nondimensional variables the constitutive law is given by eqn [5]. S ¼ K* qU=qz
½5
The nondiemensional eddy coefficient is given by eqn [6]. K* ¼ ku*0 lm =fh2m ¼ kxn
½6
Eqn [6] is the Rossby similarity relation that is the key to providing similarity solutions for stable and neutral conditions. It even provides workable results for slightly unstable conditions. Eqns [4] and [5] can be combined in an equation for nondimensionalized stress (eqn [7]). ði=K* ÞS ¼ dSd=z
½7
This has the solution eqns [8]. ˆ
S ¼ ed z
½8
dˆ ¼ ði=K* Þ1=2
½9
Eqn [8] attenuates and rotates (to the right in the Northern Hemisphere) with depth. It duplicates the salient features found in data and sophisticated numerical models.
UðzÞ Uðzsl Þ ¼
Z* z ln sl þ dˆ ðzsl zÞ k z0
for z z0 ½11
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ICE–OCEAN INTERACTION
Clearly values of the drag coefficient can vary widely depending on the under-ice roughness. Typical values of z0 range from 1 to 10 cm under pack ice. A commonly referenced value for the Arctic is 6 cm, which produces a drag coefficient at the outer edge of the log layer of 9.4 103 (Figure 1). If the reference depth is outside the log layer, the drag coefficient formulation is poorly posed because of the turning in the boundary layer. For neutral conditions, eqns [10] and [11] can be used to obtain a Rossby similarity drag law that yields the nondimensional surface drift relative to the geostrophic current for unit nondimensional surface stress (eqn [12]). U0 ¼
V0 1 ¼ ð½lnðR0 Þ A iBÞ u *0 k
½12
Here
A¼
sffiffiffiffiffiffiffiffi rffiffiffiffiffiffi! k xn þ 1 ln xn D2:2 2k 2xn sffiffiffiffiffiffiffiffi rffiffiffiffiffiffi k xn D2:3 þ B¼ 2k 2xn
salt may be expressed in kinematic form as eqns [14] and [15]. q_ ¼ /w0 T 0 S0 w0 QL
ðwith units K m s1 Þ
ðw0 þ wi ÞðS0 Sice Þ ¼ /w0 S0 S0
This Rossby similarity drag law for outside the surface layer results in a surface stress that is proportional to V1.8 rather than V2, a result that is supported by observational evidence, and can be significant at high velocities.
Heat and Mass Balance at the Ice–Ocean Interface: Wintertime Convection The energy balance at the ice–ocean interface not only exerts major influence over the ice mass balance but also dictates the seasonal evolution of upper ocean salinity and temperature structure. At low temperature, water density is controlled mainly by salinity. Salt is rejected during freezing, so that buoyancy flux from basal growth (or ablation), combined with turbulent mixing during storms, determines the depth of the well-mixed layer. Vertical motion of the ice–ocean interface depends on isostatic adjustment as the ice melts or freezes. The interface velocity is w0 þ wi where w0 ¼ ðrice =rÞh_ b ; h_ b is the basal growth rate, and wi represents isostatic adjustment to runoff of surface melt and percolation of water through the ice cover. In an infinitesimal control volume following the ice–ocean interface, conservation of heat and
½14
ðwith units psu m s1 Þ
½15 where q_ ¼ Hice =ðrcp Þ is flux (Hice) conducted away from the interface in the ice; r is water density; cp is specific heat of seawater; /w0 T0 S0 is the kinematic turbulent heat flux from the ocean; QL is the latent heat of fusion (adjusted for brine volume) divided by cp; S0 is salinity in the control volume, Sice is ice salinity, and /w0 S0 S0 is turbulent salinity flux. Fluid in the control volume is assumed to be at its freezing temperature, approximated by the freezing line (eqn [16]). T0 ¼ mS0
½13
201
½16
By standard closure, turbulent fluxes are expressed in terms of mean flow properties (eqns [17] and [18]). hw0 T 0 i0 ¼ ch u*0 dT
½17
hw0 S0 i0 ¼ cS u*0 dS
½18
u*0 is the square root of kinematic turbulent stress at the interface (friction velocity); dT ¼ T T0 and dS ¼ S S0 are differences between far-field and interface temperature and salinity; and ch and cS are turbulent exchange coefficients termed Stanton numbers. The isostatic basal melt rate, w0 is the key factor in interface thermodynamics, and in combination with wi it determines the salinity flux. A first-order approach to calculating w0 that is often sufficiently accurate (relative to uncertainties in forcing parameters) when melting or freezing is slow, is to assume that S0 ¼ S, the far-field salinity, and that ch is constant. Combining [14], [16], and [17] gives eqn [19]. w0 ¼
ch u*0 ðT þ mSÞ q_ QL
½19
Salinity flux is determined from [15]. Note the cS is not used, and that this technique fixes (unrealistically) the temperature at the interface to be the mixed layer freezing temperature. A more sophisticated approach is required when melting or freezing is intense. Manipulation of [14]
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ICE–OCEAN INTERACTION
through [18] produces a quadratic equation for w0 (eqn [20]). SL 2 w þ ðST þ SL cS Sice Þw0 þ ðu*0 cS þ wi ÞST u*0 0 þ u*0 cS S wi Sice ¼ 0 ST ¼
q_ T =m ch u*0
½20
and
SL ¼ QL =ðmch Þ ½21
Here ch and cS (turbulent Stanton numbers for heat and salt) are both important and not necessarily the same. Melting or freezing will decrease or increase S0 relative to far-field salinity, with corresponding changes in T0. The Marginal Ice Zone Experiments (MIZEX) in the 1980s showed that existing ice–ocean turbulent transfer models overestimated melt rates by a wide factor. It became clear that the rates of heat and mass transfer were less than momentum transfer (by an order of magnitude or more), and were being controlled by molecular effects in thin sublayers adjacent to the interface. If it is assumed that the extent of the sublayers is proportional to the bottom roughness scale, z0, then dimensional analysis suggests that the Stanton numbers (nondimensional heat and salinity flux) should depend mainly on two other dimensionless groups, the turbulent Reynolds number, Re* ¼ u*0 z0/n, where n is molecular viscosity, and the Prandtl (Schmidt) numbers, n/nT(S), where nT and nS are molecular diffusivities for heat and salt. Laboratory studies of heat and mass transfer over hydraulically rough surfaces suggested approximate expressions for the Stanton numbers of the form shown in eqn [22]. chðSÞ ¼
/w0 TðSÞ0 S0 n 2=3 pðRe* Þ1=2 u*0 dTðSÞ nTðSÞ
½22
The Stanton number, ch, has been determined in several turbulent heat flux studies since the original MIZEX experiment, under differing ice types with z0 values ranging from less than a millimeter (eastern Weddell Sea) to several centimeters (Greenland Sea MIZ). According to [22], ch should vary by almost a factor of 10. Instead, it is surprisingly constant, ranging from about 0.005 to 0.006, implying that the Reynolds number dependence from laboratory results cannot be extrapolated directly to sea ice. If the Prandtl number dependence of [14] holds, the ratio ch /cS ¼ (nh/nS)2/3 is approximately 30. Under conditions of rapid freezing, the solution of [20] with this ratio leads to significant supercooling of
the water column, because heat extraction far outpaces salt injection in what is called double diffusion. This result has caused some concern. Because the amount of heat represented by this supercooling is substantial, it has been hypothesized that ice may spontaneously form in the supercooled layer and drift upward in the form of frazil ice crystals. This explanation has not been supported by ice core sampling, which shows no evidence of widespread frazil ice formation beyond that at the surface of open water. The physics of the freezing process suggest that the seeming paradox of the supercooled boundary layer may be realistic without spontaneous frazil formation. When a parcel of water starts to solidify into an ice crystal, energy is released in proportion to the volume of the parcel. At large scales this manifests itself as the latent heat of fusion. However, as the parcel solidifies, energy is also required to form the surface of the solid. This surface energy penalty is proportional to the surface area of the parcel and depends on other factors including the physical character of any nucleating material. In any event, if the parcel is very small the ratio of parcel volume to surface area will be so small that the energy released as the volume solidifies is less than the energy needed to create the new solid surface. For this reason, ice crystals cannot form even in supercooled water without a nucleating site of sufficient size and suitable character. In the clean waters of the polar regions, the nearest suitable site may only be at the underside of the ice cover where the new ice can form with no nucleation barrier. Therefore, it is possible to maintain supercooled conditions in the boundary layer without frazil ice formation. Furthermore, recent results suggest that supercooling in the uppermost part of the boundary layer may be intrinsic to the ice formation process. Sea ice is a porous mixture of pure ice and high-salinity liquid water (i.e., brine). The bottom surface of a growing ice floe consists of vertically oriented pure ice platelets separated by vertical layers of concentrated brine. This platelet–brine sandwich (on edge) structure is on the scale of a fraction of a millimeter, and its formation is controlled by molecular diffusion of heat and salt. The low solid solubility of the salt in the ice lattice results in an increase of the salinity of water in the layer above the advancing freezing interface. Because heat diffuses more rapidly than salt at these scales, the cold brine tends to supercool the water below the ice–water interface. With this local supercooling, any disturbance of the ice bottom will tend to grow spontaneously. The conditions of sea ice growth are such that this
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ICE–OCEAN INTERACTION
instability is always present. Continued growth results in additional rejection of salt, some fraction of which is trapped in the brine layers, and consequently the interfacial region of the ice sheet continues to experience constitutional supercooling. Also, anisotropy in the molecular attachment efficiency intrinsic to the crystal structure of the ice platelets creates an additional supercooling in the interfacial region. The net result is that heat is extracted from the top of the water column at the rate needed to maintain its temperature near but slightly below the equilibrium freezing temperature as salt is added. This and the convective processes in the growing ice may imply that ch/cS ¼ 1 during freezing. The situation with melting may be quite different, since the physical properties of the interface change dramatically. Observations to date suggest that ch remains relatively unchanged with variable ice type and mixed layer temperature elevation above freezing. A value of 5.5 103 is representative. cS is not so well known, since direct measurements of /w0 S0 S are relatively rare. The dependence of the exchange coefficients on Prandtl and Schmidt numbers is not clear, and will only be resolved with more research.
Effects of Horizontal Inhomogeneity: Wintertime Buoyancy Flux Although the under-ice surface may be homogeneous over ice floes hundreds of meters in extent, the key fluxes of heat and salt are characteristically nonuniform. As ice drifts under the action of wind stress, the ice cover is deformed. Some areas are forced together, producing ridging and thick ice, and some areas open in long, thin cracks called leads. In special circumstances the ice may form large, unit-aspectratio openings called polynyas. In winter the openings in the ice expose the sea water directly to cold air without an intervening layer of insulating sea ice. This results in rapid freezing. As the ice forms, it rejects salt and results in unstable stratification of the boundary layer beneath open water or thin ice. These effects are so important that, even though such areas may account for less than 10% of the ice cover, they may account for over half the total ice growth and salt flux to the ocean. Thus the dominant buoyancy flux is not homogeneous but is concentrated in narrow bands or patches. Similarly, in the summer solar radiation is reflected from the ice but is nearly completely absorbed by open water. Fresh water from summertime surface melt tends to drain
203
into leads, making them sources of fresh water flux as well. The effect of wintertime convection in leads is illustrated in Figure 2. It shows two extremes in the upper ocean response. Figure 2A shows what we might expect in the case of a stationary lead. As the surface freezes, salt is rejected and forms more dense water that sinks under the lead. This sets up a circulation with fresh water flowing in from the sides near the surface and dense water flowing away from the lead at the base of the mixed layer. Figure 2B illustrates the case in which the lead is embedded in ice moving at a velocity great enough to produce a well-developed turbulent boundary layer (e.g. 0.2 m s1). If the mixed layer is fully turbulent, the cellular convection pattern may not occur; rather, the salt rejected at the surface may simply mix into the surface boundary layer. The impact of nonhomogeneous surface buoyancy flux on the boundary layer can also be characterized by the equations of motion. The viscous terms in eqn [1] can be neglected at the scales we discuss here, but the possibility of vertical motion associated with large-scale convection requires that we include the vertical component of velocity. For steady state we have eqn [23]. q qV¯ ¯ ¯ ¯ ¯ K r1 rp VdrV þ f V ¼ qz q
½23
V¯ is the velocity vector including the mean vertical velocity w; f¯ is the Coriolis parameter times the vertical unit vector. The advective acceleration term, ¯ ¯ and pressure gradient term are necessary to Vdr V, account for the horizontal inhomogeneity that is caused by the salinity flux at the lead surface. The condition that separates the free convection regime of Figure 2A and the forced convection regime of Figure 2B is expressed by the relative magnitude of the pressure gradient, r1rhp, and turbulent stress, q=qzðKqV=qzÞ, terms in [23]. This ratio can be derived with addition of mass conservation and salt conservation equations, and if we assume the vertical equation is hydrostatic, qp=qz ¼ gr ¼ gMS, where M is the sensitivity of density to salinity. If we nondimensionalize the equations by the ice velocity Ui, mixed-layer depth, d, average salt flux at the lead surface, FS, and friction velocity, u*0, the ratio of the pressure gradient term to the turbulent stress term scales as eqn [24].
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L0 ¼
gMFS d r0 Ui u2*0
½24
204
ICE–OCEAN INTERACTION
Free convection Vice ~ 0
z
Forced convection
Lead y
Lead
LL
Vice u*
x
FS
du (x)
u (z )
Thick sea ice dd (x)
Increased S dml u (z )
Internal boundary layers, depth = d (x )
(z )
(A)
(B)
Figure 2 Modes of lead convection. (A) The free convection pattern that results when freezing and salt flux are strong, and the relative velocity of the ice is low. Cellular patterns of convective overturning are driven by pressure gradients that arise from the salinity distrurbance due to ice formation. (B) The forced convection regime that exists when ice motion is strong. The salinity flux and change in surface stress in the lead cause a change in the character of the boundary layer that grows deeper downstream. The balance of forces is primarily Coriolis and turbulent diffusion of momentum. (From Morison JH, MCPhee MG, Curtin T and Paulson CA (1992) The oceanography of winter leads. Journal of Geophysical Research 97: 11199–11218.)
autonomous underwater vehicle. Using the vehicle vertical motion as a proxy for vertical water velocity, it is also possible to estimate the salt flux w0 S0 . The lead was moving at 0.04 m s1, and estimates of salt flux put L0 between 4 and 11 (free convection in Figure 3). Salinity increased in the downstream direction across the lead and reached a sharp maximum Water temperature _ air temperature (˚C)
If this lead number is small because the ice is moving rapidly or the salt flux is small, the pressure gradient term is not significant in [23]. In this forced convection case, illustrated in Figure 2B, the boundary layer behaves as in the horizontally homogeneous case except that salt is advected and diffused away from the lead in the turbulent boundary layer. If the lead number is large because the ice is moving slowly or the salt flux is large, the pressure gradient term is significant. In this free convention case the salinity disturbance is not advected away, but builds up under the lead. This creates pressure imbalances that can drive the type of cellular motion shown in Figure 2A. Figure 3 shows conditions for which the lead number is unity for a range of ice thickness. Here the salt flux has been parametrized in terms of the air– sea temperature difference, and stress has been parametrized in terms of Ui. The figure shows the locus of points where L0 is equal to unity. For typical winter and spring conditions, L0 is close to 1, indicating that a mix of free and forced convection is common. Conditions where lead convection features have been observed are also shown in Figure 3. Most of these are in the free convection regime, probably because they are more obvious during quiet conditions. There have been several dedicated efforts to study the effects of wintertime lead convection. The most recent example was the 1992 Lead Experiment (LeadEx) in the Beaufort Sea. Figure 4 illustrates the average salinity profile at 9 m under a nearly stationary lead. The data was gathered with an
30 20 15 10 hi = 5 cm
′71
25 A3
20
A4
′92 Lead 3
′92 Lead 4
h i = 0 cm A
15 Free convection
10
Forced convection
5
L0 = 1 L0 t = 1
0 0
0.02
0.04 0.06 0.08
0.10 0.12
0.14 0.16
_
Ice velocity (m s 1) Figure 3 Air–water temperature difference versus Ui for L0 equal to 1 for various ice thicknesses, hi. Also shown are the temperature difference and ice velocity values for several observations of lead convection features such as underice plumes. Most of these are in the free convection regime: 0 71 denotes the AIDJEX pilot study; A3 denotes the 1974 AIDJEX Lead Experiment – lead 3 (ALEX3); A4 denotes ALEX4; A denotes the 1976 Arctic Mixed Layer Experiment; and 0 92 Lead 4 denotes the 1992 LeadEx lead 4. LeadEx lead 3 (0 92 Lead 3) was close to L0 ¼ 1. (From Morison JH, McPhee MG, Curtin T and Paulson CA (1992) The oceanography of winter leads. Journal of Geophysical Research 97: 11199–11218.)
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ICE–OCEAN INTERACTION
205
ICE
ICE
LEAD
2
0
Relative current
_
Salt flux down (10 5 kg m 2 s 1)
_3
Salinity fluctuation (10 PSU)
4
_
_
~0.02−0.04 m s
_1
Lead average
5
= 6.0 × 10
_6
0
_ 250
_ 200
_ 150
_ 100
_ 50
0
50
100
150
200
Cross-lead distance (m) Figure 4 Composites of S 0 and w 0 S 0 at 9 m depth measured with an autonomous underwater vehicle during four runs under lead 4 at the 1992 Lead Experiment. The horizontal profile data have been collected in 1 m bins. (From Morison JH, McPhee MG (1998) Lead convection measured with an autonomous underwater vechicle, Journal of Geophysical Research 103: 3257–3281.)
at the downstream edge. The salt flux was highest near the lead edges, but particularly at the downstream edge. With even a slight current, the downstream edge plume is enhanced by several factors. The vorticity in the boundary layer reinforces the horizontal density gradient at the downstream edge and counters the gradient at the upstream edge. The salt excess is greatest at the downstream edge by virtue of the salt that is advected from the upstream lead surface. The downstream edge plume is also enhanced by the vertical motion of water at the surface due to water the horizontal flow being forced downward under the ice edge. Figure 5 shows the salt flux beneath a 1000 m wide lead moving at 0.14 m s1 with L0 equal to about 1 (Figure 3). Here the salt flux is more evenly spread under the lead surface. The salt flux derived from the direct w0 S0 correlation method does show some enhancement at the lead edge. This may be partly due to the influence of pressure gradient forces and the reasons cited for the free convection case described above. The other factor that influences the convective pattern is the lead width. In the case of the 100 m lead in even a weak current, the convection may not be fully developed until the downstream edge is reached. For the 1000 m lead of the second case, the convection under the downstream portion of the
lead was a fully developed unstable boundary layer. The energy-containing eddies filled the mixed layer and their dominant horizontal wavelength was equal to about twice the mixed layer depth.
Effects of Horizontal Inhomogeneity: Summertime Buoyancy Flux The behavior of the boundary layer under summer leads is relatively unknown compared to the winter lead process. Because of the important climate consequences, it is a subject of increasing interest. Summertime leads are thought to exhibit a critical climate-related feature of air–sea–ice interaction, icealbedo feedback. This is because leads are windows that allow solar radiation to enter the ocean. The proportion of radiation that is reflected (albedo) from sea ice and snow is high (0.6–0.9) while that from open water is low (0.1). The fate of the heat that enters summer leads is important. If it penetrates below the draft of the ice, it warms the boundary layer and is available to melt the bottom of the ice over a large area. If most of the heat is trapped in the lead above the draft of the ice, it will be available to melt small pieces of ice and the ice
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ICE–OCEAN INTERACTION
30
LEAD
_6
_2
_1
Salt flux (10 kg m s )
25 20 Salt flux from Ebp-IDM
15 10
Lead average flux _6 7.762 × 10
Downstream average _ flux = 3.362 × 10 6
5 0 w ′S ′ salt flux, 56-m bin composite
_5
−6
average = 9.41× 10 , lead average = 1.19 ×10
−5
_ 10 _ 300
_ 200
_ 100
0
100
200
300
400
Distance upstream from the lead edge (m) Figure 5 Composite average for autonomous underwater vehicle runs 1 to 5 at lead 3 of the 1992 Lead Experiment. The salinity is band-passed at 1 rad m1 and is indicative of the turbulence level and is used to estimate the salt flux by the Ebp IDM method of Morison and McPee (1998). The salt flux is elevated in the lead and decreases beyond about 72 m downstream of the lead edge. The composite average of w 0 S 0 in 56 m bins for the same runs is also shown in the center panel. The average flux and the decrease downstream are about the same as given by the Ebp IDM method, but w 0 S 0 suggests elevated fluxes near the lead edge. (From Morison JH, McPhee MG (1998) Lead convection measured with an autonomous underwater vechicle, Journal of Geophysical Research 103: 3257–3281.)
floe edges. In the latter case the area of ice will be reduced and the area of open water increased. This allows even more solar radiation to enter the upper ocean, resulting in a positive feedback. This process may greatly affect the energy balance of an ice-covered sea. The critical unknown is the partition of heating between lateral melt of the floe edges and bottom melt. There are fundamental similarities between the summertime and wintertime lead problems. The equations of motion ([15]–[24]) are virtually identical. Only the sign of the buoyancy flux is opposite. The heat flux is important to summer leads and tends to decrease the density of the surface waters. However, as with winter leads, the buoyancy flux is controlled mainly by salt. As the top surface of the ice melts, much of the water that does not collect in melt ponds on the ice surface instead runs into the leads. If the ambient ice velocity is low, ice melt from the bottom surface will tend to flow upward and collect in the leads as well. Thus leads are the site of a concentrated flux of fresh water accumulated over large areas of ice. If this flux, FS, into the lead is negative enough relative to the momentum flux represented by u*0, the lead number, L0, will be a large negative number and shear production of turbulent energy will not be able to overcome the stabilizing buoyant production. This means turbulent mixing will be weak beneath the lead surface and a layer of fresh water will accumulate near the surface of the lead. The stratification
at the bottom of this fresh water layer may be strong enough to prevent mixing until a storm produces a substantial stress. This will be made even more difficult than in the winter situation because of the effect of stabilizing buoyancy flux on the boundary layer generally. The only way the fresh water will be mixed downward is by forced convection; there is no analogue to the wintertime free convection regime. When there is sufficient stress to mix out a summertime lead, the pattern must resemble that of the forced convection regime in Figure 2A. At the upstream edge of the lead, fresh warm water will be mixed downward in an internal boundary layer that increases in thickness downstream until it reaches the steady-state boundary layer thickness appropriate for that buoyancy flux or the ambient mixed layer depth. The rate of growth should scale with the local value of u*0 (or perhaps u*0 Z* ). At the downstream edge, another boundary layer conforming to the under-ice buoyancy flux and surface stress will begin to grow at a rate roughly scaling with the local u*0. In spite of the generally stabilizing buoyancy flux, this process has the effect of placing colder, more saline water from under the ice on top of fresher and warmer (consequently lighter) water drawn from the lead. Thus, even embedded in the stable summer boundary layer, the horizontal inhomogeneity due to leads may create pockets of instability and more rapid mixing than might be expected on the basis of average conditions.
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ICE–OCEAN INTERACTION
Recent studies of summertime lead convection at the 1997–98 Surface Heat Budget of the Arctic experiment saw the salinity decrease in the upper 1 m of leads to near zero and temperatures increase to more than 01C. Only when ice velocities were driven by the wind to speeds of nearly 0.2 m s1 were these layers broken down and the fresh, warm water mixed into the upper ocean. At these times the heat flux measured at 5 m depth reached values over 100 W m2. The criteria for the onset of mixing are being studied along with the net effect of the growing internal boundary layers. Even with an understanding of the mixing process, it will be a challenge to apply this information to larger-scale models, because the mixing is nonlinearly dependent on the history of calm periods and strong radiation.
Internal Waves and Their Interaction with the Ice Cover One of the first studies of internal waves originated with observations made by Nansen during his 1883 expedition. It did not actually involve interaction with the ice cover, but with his ship the Fram. He found that while cruising areas of the Siberian shelf covered with a thin layer of brackish water, the Fram had great difficulty making any headway. It was hypothesized by V. Bjerknes and proved by Ekman that this ‘dead water’ phenomenon was caused by the drag of the internal wave wake produced by the ship’s hull as it passed through the shallow surface layer. This suggests that internal wave generation by deep keels may cause drag on moving ice. Evidence of internal wave generation by keels has been observed by several authors, but estimates of the amount of drag vary widely. This is due mainly to wide differences in the separation of the stratified pycnocline and the keels. The drag produced by under-ice roughness of amplitude h0 with horizontal wavenumber b moving at velocity Vi (magnitude vi) over a pycnocline with stratification given by Brunt–Vaisala frequency, N, a depth d below the ice–ocean interface, can be expressed as an effective internal wave stress (eqn [25]), where Cwd (eqn [26]) accounts for the drag that would exist if there were no mixed layer between the ice and the pycnocline. Siw ¼ GCwd Vi
½25
above which the waves are evanescent (bc ¼ N/vi). G is an attenuation factor that accounts for the separation of the pycnocline from the ice by the mixed layer of depth d (eqn [27]). 8" 911 #2 < = 2 bDb N G ¼ @sinh2 ðbdÞ cothðbdÞ 2 2 þ 2 2 1 A : ; ui b x ui bx 0
½27 Db is the strength of the buoyancy jump at the base of the mixed layer. For wavenumbers of interest and d much bigger than about 10 m, G becomes small and internal wave drag is negligible. Thus it is not a factor in the central Arctic over most of the year. However, in the summer pack ice, and many times in the marginal ice zone, stratification will extend to or close to the surface. Then internal wave drag can be at least as important as form drag. The ice cover also uniquely affects the ambient internal wave field. In most of the world ocean the internal wave energy level, when normalized for stratification, is remarkably uniform. It has been established by numerous studies that the internal wave energy in the Arctic Ocean is typically several times lower. In part this may be due to the absence of surface gravity waves. The other likely reason is that friction on the underside of the ice damps internal waves. Decomposing the internal wave field into vertical modes, one finds the mode shapes for horizontal velocity are a maximum at the surface. This is perfectly acceptable in the open water situation. However, at the horizontal scales of most internal waves, an ice cover imposes a surface boundary condition of zero horizontal velocity. The effect of this can be estimated by assuming that a time-varying boundary layer is associated with each spectral component of the internal wave field. This is not rigorously correct because all the modes interact in the same nonlinear boundary layer, and are thereby coupled. However, in the presence of a dominant, steady current due to ice motion, the effect on the internal wave modes can be linearized and considered separately. The near-surface internal wave velocity can be approximated as a sum of rotary components (eqn [28]). VðzÞ ¼
M P
Dn ðzÞeion t ¼
n¼0
Cwd ¼ 12b2x h0 ½ðb2c =b2x Þ 11=2
M P
½An ðzÞ þ iBn ðzÞeion t
n¼0
½26
½28
The wavenumber in the direction of the relative ice velocity, Vi, is bx, and bc is the critical wave number
The internal wave motion away from the boundary DNn can be subtracted from the linear
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ICE–OCEAN INTERACTION
time-varying boundary layer equation (eqn [1] with the addition of the time variation acceleration, qV/qt). This yields an equation for each rotary component of velocity in the boundary layer (eqn [29]). iðon þ f ÞðDn DNn Þ ¼
Dn ¼ 0 Dn ¼ DNn
q qDn K qz qz
½29
at z ¼ z0 at z ¼ d
This oscillating boundary layer equation can be solved for K of the form K ¼ ku*0 zexp ( 6| þ f |zu*0 ). When we do this for representative internal wave conditions in the Arctic and compute the energy dissipation, we find the timescale required to dissipate the internal wave energy through underice friction is 32 days. This is a factor of 3 smaller than is typical for open ocean conditions. Assuming a steady state with internal wave forcing and other dissipation mechanisms in place, the under-ice boundary layer will result in a 75% reduction in steady-state internal wave energy. This suggests the effect of the under-ice boundary layer is critical to the unique character of internal waves in ice-covered seas.
Outstanding Issues The outstanding issue of ice–ocean interaction is how the small-scale processes in the ice and at the interface affect the exchange between the ice and water. This is arguably most urgent in the case of heat and salt exchange during ice growth. When we apply laboratory-derived concepts for the diffusion of heat and salt to the ice–ocean interface, we get results that are not supported by observation, such as spontaneous frazil ice formation and large ocean heat flux under thin ice. These results are causing significant errors in largescale models. They stem from a molecular sublayer model of the ice–ocean interface (Figure 1) and the difference between the molecular diffusivities of heat and salt. What seems to be wrong is the molecular sublayer model. Recent results in the microphysics of ice growth reveal that the structure and thermodynamics of the growing ice produce instabilities and convection within the ice and extending into the water. The ice surface is thus not a passive, smooth surface covered with a thin molecular layer. Rather it is field of jets emitting plumes of supercooled, high-salinity water at a very small scale. This type of
unstable convection likely tends to equalize the diffusion of heat and salt relative to the apparently unrealistic parameterizations we are using now. Similarly, we do not really understand how the turbulent stress we might measure in the surface layer is converted to drag on the ice. Certainly a portion of this is through viscous friction in the molecular sublayer. However, in most cases the underside of the ice is not hydrodynamically smooth, which suggests that pressure force acting on the bottom roughness elements are ultimately transferring a large share of the momentum. Understanding this will require perceptual breakthroughs in our view of how turbulence and the mean flow interact with a rough surface buried in a boundary layer. Achieving this understanding is complicated greatly by a lack of contemporaneous measurements of turbulence and under-ice topography at the appropriate scales. This drag partition problem is general and not limited to the under-ice boundary layer. However, the marvelous laboratory that the under-ice boundary layer provides may be the place to solve it.
See also Coupled Sea Ice–Ocean Models. Internal Waves. Sea Ice: Overview. Sub Ice-Shelf Circulation and Processes. Under-Ice Boundary Layer.
Further Reading Johannessen OM, Muench RD, and Overland JE (eds.) (1994) The Polar Oceans and Their Role in Shaping the Global Environment: The Nansen Centennial Volume. Washington, DC: American Geophysical Union. McPhee MG (1994) On the turbulent mixing length in the oceanic boundary layer. Journal of Physical Oceanography 24: 2014--2031. Morison JH, McPhee MG, and Maykutt GA (1987) Boundary layer, upper ocean and ice observations in the Greenland Sea marginal ice zone. Journal of Geophysical Research 92(C7): 6987--7011. Morison JH and McPhee MG (1998) Lead convection measured with an autonomous underwater vehicle. Journal of Geophysical Research 103(C2): 3257--3281. Smith WO (ed.) (1990) Polar Oceanography. San Diego, CA: Academic Press. Wettlaofer JS (1999) Ice surfaces: macroscopic effects of microscopic structure. Philosphical Transactions of the Royal Society of London A 357: 3403--3425.
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ICE SHELF STABILITY C. S. M. Doake, British Antarctic Survey, Cambridge, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1282–1290, & 2001, Elsevier Ltd.
Introduction Ice shelves are floating ice sheets and are found mainly around the Antarctic continent (Figure 1). They can range in size up to 500 000 km2 and in thickness up to 2000 m. Most are fed by ice streams and outlet glaciers, but some are formed by icebergs welded together by sea ice and surface accumulation. They exist in embayments where the shape of the bay and the presence of ice rises, or pinning points, plays an important role in their stability. Ice shelves lose mass by basal melting, which modifies the properties of the underlying water mass and eventually influences the circulation of the global ocean, and by calving. Icebergs, sometimes more than 100 km in length, break off intermittently from the continent and drift to lower latitudes, usually breaking up and melting by the time they reach the Antarctic Convergence. The lowest latitude that ice shelves can exist is determined by the mean annual air temperature. In the Antarctic Peninsula a critical isotherm of about 51C seems to represent the limit of viability. In the last 40 years or so, several ice shelves have disintegrated in response to a measured atmospheric warming trend in the western and northern part of the Antarctic Peninsula. However, this warming trend is not expected to affect the stability of the larger ice shelves further south such as Filchner–Ronne or Ross, in the near future.
What are Ice Shelves? Physical and Geographical Setting
An ice shelf is a floating ice sheet, attached to land where ice is grounded along the coastline. Nourished mainly by glaciers and ice streams flowing off the land, ice shelves are distinct from sea ice, which is formed by freezing of sea water. Most ice shelves are found in Antarctica where the largest can cover areas of 500 000 km2 (e.g. Ross Ice Shelf and Filchner–Ronne Ice Shelf). Thicknesses vary from nearly 2000 m, around for example parts of the grounding line of Ronne Ice Shelf, to about 100 m at the seaward edge known as the ice front.
Typically, ice shelves exist in embayments, constrained by side walls until they diverge too much for the ice to remain in contact. Thus the geometry of the coastline is important for determining both where an ice shelf will exist and the position of the ice front. There are often localized grounding points on sea bed shoals, forming ice rises and ice rumples, both in the interior of the ice shelf and along the ice front. These pinning points provide restraint and cause the ice shelf to be thicker than if it were not pinned. Ice flows from the land to the ice front. Input velocities range from near zero at a shear margin (e.g., with a land boundary), to several hundred meters per year at the grounding line where ice streams and outlet glaciers enter. At the ice front velocities can reach up to several kilometers per year. A characteristic of ice shelves is that the (horizontal) velocity is almost the same at all depths, whereas in glaciers and grounded ice sheets the velocity decreases with depth (Figure 2). In the Antarctic, most ice shelves have net surface accumulation although there may be intensive summer melt which floods the surface. The basal regime is controlled by the subice circulation. Basal melting is often high near both the grounding line and the ice front. Marine ice can accumulate in the intermediate areas, where water at its in situ freezing point upwells and produces frazil ice crystals. Surface temperatures will be near the mean annual air temperature whereas the basal temperature will be at the freezing point of the water. Therefore the coldest ice is normally in the upper layers. Ice shelves normally form where ice flows smoothly off the land as ice streams or outlet glaciers. In some areas, however, ice breaks off at the coastline and reforms as icebergs welded together by frozen sea ice and surface accumulation. The processes of formation and decay are likely to operate under very different conditions. An ice shelf is unlikely to reform under the same climatic conditions which caused it to decay. Complete collapse is a catastrophic process and rebuilding requires a major change in the controlling parameters. However, there is evidence of cyclicity on periods of a few hundred years in some areas. Collapse of the northernmost section of Larsen Ice Shelf within a few days in January 1995 indicates that, after retreat beyond a critical limit, ice shelves can disintegrate rapidly. The breakup history of two northern sections of Larsen Ice Shelf (Larsen A and Larsen B) between 1986 and 1997 has been used to determine a stability criterion for ice shelves.
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Figure 1 Distribution of ice shelves around the coast of Antarctica.
Figure 2 Cartoon of ice shelf in bay with ice streams.
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Analysis of various ice-shelf configurations reveals characteristic patterns in the strain-rates near the ice front which have been used to describe the stability of the ice shelf.
Why are Ice Shelves Important? Ice shelves are one of the most active parts of the ice sheet system. They interact with the inland ice sheet, glaciers and ice streams which flow into them, and with the sea into which they eventually melt, either directly or as icebergs. A typical Antarctic ice shelf will steadily advance until its ice front undergoes periodic calving, generating icebergs. Calving can occur over a wide range of time and spatial scales. Most ice fronts will experience a quasi-continuous ‘nibbling’ away of the order of tens to hundreds of meters per year, where ice cliffs collapse to form bergy bits and brash ice. The largest icebergs may be more than 100 km in size, but will calve only at intervals of 50 years or more. Thus when charting the size and behavior of an ice shelf it can be difficult to separate and identify stable changes from unstable ones. Mass balance calculations show that calving of icebergs is the largest factor in the attrition of the Antarctic Ice Sheet. Estimates based on both ship and satellite data suggest that iceberg calving is only slightly less than the total annual accumulation. Ice shelf melting at the base is the other principal element in the attrition of the ice sheet, with approximately 80% of all ice shelf melting occurring at distances greater than 100 km from the ice front. Although most of the mass lost in Antarctica is through ice shelves, there is still uncertainty about the role ice shelves play in regulating flow off the land, which is the important component for sea-level changes. Because ice shelves are already floating, they do not affect sea level when they breakup. Ice shelves are sensitive indicators of climate change. Those around the Antarctic Peninsula have shown a pattern of gradual retreat since about 1950, associated with a regional atmospheric warming and increased summer surface melt. The effects of climate change on the mass balance of the Antarctic Ice Sheet and hence global sea level are unclear. There have been no noticeable changes in the inland ice sheet from the collapse of the Antarctic Peninsula ice shelves. This is because most of the margins with the inland ice sheet form a sharp transition zone, making the ice sheet dynamics independent of the state of the ice shelf. This is not the case with a grounding line on an ice stream where there is a smooth transition zone, but generally the importance of the local stresses there diminishes rapidly upstream.
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An important role of ice shelves in the climate system is that subice-shelf freezing and melting processes influence the formation of Antarctic Bottom Water, which in part helps to drive the oceanic thermohaline circulation (see Sub Ice-Shelf Circulation and Processes).
How have Ice Shelves Changed in the Past? As the polar ice sheets have waxed and waned during ice age cycles, ice shelves have also grown and retreated. During the last 20–25 million years, the Antarctic Ice Sheet has probably been grounded out to the edge of the continental shelf many times. It is unlikely that any substantial ice shelves could have existed then, because of unsuitable coastline geometry and lack of pinning points, although the extent of the sea ice may have been double the present day area. At the last glacial maximum, about 20 000 years ago, grounded ice extended out to the edge of the continental shelf in places, for example in Prydz Bay, but not in others such as the Ross Sea. When the ice sheets began to retreat, some of the large ice shelves seen today probably formed by the thinning and eventual flotation of the formerly grounded ice sheets. Large ice shelves also existed in the Arctic during the Pleistocene. Abundant geologic evidence shows that marine Northern Hemisphere ice sheets disappeared catastrophically during the climatic transition to the current interglacial (warm period). Only a few small ice shelves and tidewater glaciers exist there today, in places like Greenland, Svalbard, Ellesmere Island and Alaska. There have been periodic changes in Antarctic iceshelf grounding lines and ice fronts during the Holocene. The main deglaciation on the western side of the Antarctic Peninsula which started more than 11 000 years ago was initially very rapid across a wide continental shelf. By 6000 years ago, the ice sheet had cleared the inner shelf, whereas in the Ross Sea retreat of ice shelves ceased about the same time. Retreat after the last glacial maximum was followed by a readvance of the grounding line during the climate warming between 7000 and 4000 years ago. Open marine deposition on the continental shelf is restricted to the last 4000 years or so. Short-term cycles (every few hundred years) and longer-term events (approximately 2500 year cycles) have been detected in marine sediment cores that are likely related to global climate fluctuations. More recently, ice-shelf disintegration around the Antarctic Peninsula has been associated with a regional atmospheric warming which has been occurring
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since about 1950. Ice front retreat of marginal ice shelves elsewhere in Antarctica (e.g., Cook Ice Shelf, West Ice Shelf) is probably also related to atmospheric warming. There is little sign of any significant impact on the grounded ice.
Where are Ice Shelves Disintegrating Now? Two Case Studies The detailed history of the breakup of two ice shelves, Wordie and Larsen, illustrate some of the critical features of ice-shelf decay. Wordie Ice Shelf
Retreat of Wordie Ice Shelf on the west coast of the Antarctic Peninsula has been documented using high resolution visible satellite images taken since 1974 and the position of the ice front in 1966 mapped from aerial photography. First seen in 1936, the ice front has fluctuated in position but with a sustained retreat starting around 1966 (Figure 3). The ice shelf area has decreased from about 2000 km2 in 1966 to about 700 km2 in 1989. However, defining the position of the ice front can be very uncertain, with large blocks calving off to form icebergs being difficult to
classify as being either attached to, or separated from, the ice shelf. Ice front retreat until about 1979 occurred mainly by transverse rifting along the ice front, creating icebergs up to 10 km by 1 km. The western area was the first to be lost, between 1966 and 1974. Results from airborne radio-echo sounding between 1966 and 1970 suggested that the ice shelf could be divided into a crevassed eastern part and a rifted western part where brine could well-up and infiltrate the ice at sea level. By 1974 there were many transverse rifts south of Napier Ice Rise and by 1979 longitudinal rifting was predominant. Many longitudinal rifts had formed upstream of several ice rises by 1979, some following preexisting flowline features. A critical factor in the break-up was the decoupling of Buffer Ice Rise; by 1986 ice was streaming past it apparently unhindered, in contrast to the compressive upstream folding seen in earlier images. Between 1988 and 1989 the central part of the ice shelf, consisting of broken ice in the lee of Mount Balfour, was lost, exposing the coastline and effectively dividing the ice shelf in two. By 1989, longitudinal rifting has penetrated to the grounding line north of Mount Balfour. Further north, a growing shear zone marked the boundary between fast-flowing ice from the north Forster Ice Piedmont and slower moving
Figure 3 Cartoon showing retreat of Wordie ice front between 1966 and 1989.
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ice from Hariot Glacier. There has been no discernible change in the grounding line position. The major ice rises have played several roles in controlling ice-shelf behavior. When embedded in the ice shelf, they created broken wakes downstream, and zones of compression upstream which helped to stabilize the ice shelf. During ice front retreat, they temporarily pinned the local ice front position and also acted as nucleating points for rifting which quickly stretched upstream, suggesting that a critical fracture criterion had been exceeded. At this stage, an ice rise, instead of protecting the ice shelf against decay, aided its destruction by acting as an indenting wedge. Breakup was probably triggered by a climatic warming which increased ablation and the amount of melt water. Laboratory experiments show that the fracture toughness of ice is reduced at higher temperatures and possibly by the presence of water. Instead of refreezing in the upper layers of firn, free water could percolate down into crevasses and, by increasing the pressure at the bottom, allow them to grow into rifts or possibly to join up with basal crevasses. Processes like these would increase the production rate of blocks above that required for a ‘steady state’ ice front position. The blocks will drift away as icebergs if conditions, such as bay geometry and lack of sea ice, are favorable. Thus, ice front retreat would be, inter alia, a sensitive function of mean annual air temperature. Some ice shelves, such as Brunt, are formed from blocks that break off at the coast line and, unable to float away, are ‘glued’ together by sea ice and snowfall. These heterogeneous ice shelves contrast with those, such as Ronne, where glaciers or ice streams flow unbroken across the grounding line to form a more homogeneous type of ice shelf. Before 1989, Wordie Ice Shelf consisted of a mixture of both types, the main tongues being derived from glacier inputs, while the central portion was formed from blocks breaking off at the grounding line and at the sides of the main tongues. The western rifted area that broke away between 1966 and 1974 was described in early 1967 as ‘snowed-under icebergs’ and it was the heterogeneous central part that broke back to the coastline around Mount Balfour in 1988/89. Increased ablation would not only enhance rifting along lines of weakness but would also loosen the ‘glue’ that held the blocks together. Larsen Ice Shelf
The most northerly ice shelf in the Antarctic, Larsen Ice Shelf, extends in a ribbon down the east coast of the Antarctic Peninsula from James Ross Island to
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the Ronne Ice Shelf. It consists of several distinct ice shelves, separated by headlands. The ice shelf in Prince Gustav Channel, between James Ross Island and the mainland, separated from the main part of Larsen Ice Shelf in the late 1950s and finally disappeared by 1995. The section between Sobral Peninsula and Robertson Island, known as Larsen A, underwent a catastrophic collapse at the end of January 1995 when it disintegrated into a tongue of small icebergs, bergy bits and brash ice. Previously the ice front had been retreating for a number of years, but in a more controlled fashion by iceberg calving. Before the final breakup, the surface that had once been flat and smooth had become undulating, suggesting that rifting completely through the ice had occurred. Collapse during a period of intense north-westerly winds and high temperatures was probably aided by a lack of sea ice, allowing ocean swell to penetrate and add its power to increasing the disintegration processes. The speed of the collapse and the small size of the fragments of ice (the largest icebergs were less than 1 km in size) implicate fracture as the dominant process in the disintegration (Figure 4). The ice front of the ice shelf known as Larsen B, between Robertson Island and Jason Peninsula, steadily advanced for about 6 km from 1975 until 1992. Small icebergs broke away from the heavily rifted zone south of Robertson Island after July 1992 and a major rift about 25 km in length had opened up by then. This rift formed the calving front when an iceberg covering 1720 km2 and smaller pieces corresponding to a former ice shelf area of 550 km2 broke away between 25 and 30 January 1995, coincident with the disintegration of Larsen A. The ice front has retreated continuously by a few kilometers per year since then (to 1999) and the ice shelf is considered to be under threat of disappearing completely.
Why do Ice Shelves Break-up? Calving and Fracture
Calving is one of the most obvious processes involved in ice shelf breakup. Although widespread, occurring on grounded and floating glaciers, including ice shelves, as well as glaciers ending in freshwater lakes, it is not well understood. The basic physics, tensile propagation of fractures, may be the same in all cases but the predictability of calving may be quite different, depending on the stress field, the basal boundary conditions, the amount of surface water, etc. Fracture of ice has been studied mainly in laboratories and applying these results to the
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Figure 4 Photo of the Larsen breakup. Picture taken approximately 10 km north of Robertson Island, looking west north-west. In the far background the mountains of the Antarctic Peninsula can be seen.
conditions experienced in naturally occurring ice masses requires extrapolations which may not be valid. Questions such as how is crevasse propagation influenced by the ice shelf geometry and by the physical properties of the ice such as inhomogeneity, crystal anisotropy, temperature and presence of water, need to be answered before realistic models of the calving process can be developed. The engineering concepts of fracture mechanics and fracture toughness promise a way forward. Fracture of ice is a critical process in ice shelf dynamics. Mathematical models of ice shelf behavior based on continuum mechanics, which treat ice as a nonlinear viscous fluid will have to incorporate fracture mechanics to describe iceberg calving and how ice rises can initiate fracture both upstream and downstream. High stresses generated at shear margins around ice rises, or at the ‘corner points’ of the ice shelf where the two ends of the ice front meet land, can exceed a critical stress and initiate crevassing. These crevasses will propagate under a favorable stress regime, and if there is sufficient surface water may rift through the complete thickness. Transverse crevasses parallel to the ice front act as sites for the initiation of rifting and form lines where calving may eventually occur. Calving is a very efficient method of getting rid of ice from an ice shelf – once an iceberg has formed, it can drift away within a few days. Sometimes,
however, an iceberg will ground on a sea bed shoal, perhaps for several years. Once an Antarctic iceberg has drifted north of the Antarctic Convergence, it usually breaks-up very quickly in the warmer waters. Climate Warming
Circumstantial evidence links the retreat of ice shelves around the Antarctic Peninsula to a warming trend in atmospheric temperatures. Observations at meteorological stations in the region show a rise of about 2.51C in 50 years. Ice shelves appear to exist up to a climatic limit, taken to be the mean annual 51C isotherm, which represents the thermal limit of ice-shelf viability. The steady southward migration of this isotherm has coincided with the pattern of iceshelf disintegration. There are too few sea temperature measurements to show if the observed atmospheric warming has resulted in warmer waters, and thus increased basal melting. Although basal melting may have increased, the indications are that it is surface processes that have played the dominant role in causing breakup. The mechanism for connecting climate warming with ice-shelf retreat is not fully understood, but increased summer melting producing substantial amounts of water obviously plays a part in enhancing fracture processes, leading to calving.
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ICE SHELF STABILITY
The retreat of ice shelves around the Antarctic Peninsula that has occurred in the last half of the twentieth century has raised questions about whether or not this reflects global warming or whether it is only a regional phenomenon. The warming trend seems to be localized to the Antarctic Peninsula region and there is no significant correlation with temperature changes in the rest of Antarctica. Stress Patterns
Numerical models of ice shelves have been used to examine their behavior and stability criteria. The strain-rate field can be specified by the principal values (e_1 and e_2 where e_1 4e_2 ) and the direction of the principal axes. A simple representation is given by the trajectories, which are a set of orthogonal curves whose directions at any point are the directions of the principal axes. Analyses of the strainrate trajectories for Filchner Ronne Ice Shelf and for different ice shelf configurations of Larsen Ice Shelf show characteristic patterns of a ‘compressive arch’ and of isotropic points (Figure 5). The ‘compressive arch’ is seen in the pattern formed by the smallest principal component (e_2 ) of the strain-rate trajectories. Seaward of the arch both principal strain-rate components are extensive, whereas inland of the arch the e_2 component is compressive. The arch extends from the two ends of the ice front across the whole width of the ice shelf. It is a generic feature of ice shelves studied so far and structurally stable to small perturbations. It is probably related to the geometry of the ice shelf bay. A critical arch, consisting entirely of compressive trajectories, appears to correspond to a criterion for stability. If the ice front breaks back through the arch then an irreversible retreat occurs, possibly catastrophically, to another stable configuration. The exact location of the critical arch cannot yet be determined a priori, but it is probably close to the compressive arch delineated by the transition from extension to compression for e_2. Another pattern seen in the strain-rate trajectories is that of isotropic points. They are indicators of generic features in the surface flow field which are stable to small perturbations in the flow. This means that their existence should not be sensitive either to reasonable errors in data used in the model or to simplifications in the model itself. They act as (permanent) markers in a complicated (tensor) strainrate field and are often located close to points where the two principal strain-rates are equal or where the velocity field is stationary (usually either a maximum or a saddle point). Isotropic points are classified by a number of properties, but only two categories can be
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reliably distinguished observationally, by the way the trajectory of either of the principal strain-rates varies around a path enclosing the isotropic point. In one case the trajectory varies in a prograde sense and the isotropic point is called a ‘monstar’, whereas in the other case the trajectory varies in a retrograde sense and the isotropic point is called a ‘star’ (Figure 5). The two categories of isotropic points can be identified in the model strain-rate trajectories, ‘stars’ occurring near input glaciers and ‘monstars’ near ice fronts if they are in a stable configuration. Icebergs calving off Filchner Ice Shelf in 1986 moved the position of the ‘monstar’ up to the newly formed ice front, whereas on Ronne Ice Shelf the ‘monstar’ is about 50 km inland of the ice front. This suggests that the existence of a monstar can be used as a ‘weak’ indicator of a stable ice front. Calving can remove a monstar even if the subsequent ice front position is not necessarily an unstable one and may readvance.
How Vulnerable are the Large Ice Shelves and how Stable are Grounding Lines? The two largest ice shelves (Ross and Filchner– Ronne) are too far south to be attacked by the atmospheric warming that is predicted for the twentyfirst century. Basal melting will increase if warmer water intrudes onto the continental shelf, but the effects of a warming trend could be counterintuitive. If the warming was sufficient to reduce the rate of sea ice formation in the Weddell Sea, then the production of High Salinity Shelf Water (HSSW) would reduce as well. This would affect the subice shelf circulation, replacing the relatively warm HSSW under the Filchner–Ronne Ice Shelf with a colder Ice Shelf Water which would reduce the melting and thus thicken the ice shelf. Further climate warming would restore warmer waters and eventually thin the ice shelf by increasing the melting rates, possibly to values of around 15 m year 1 as seen under Pine Island Glacier. The likelihood of this is discussed in Sub Ice-Shelf Circulation and Processes. The supplementary question is what effect, if any, would the collapse of the ice shelves have on the ice sheet, especially the West Antarctic Ice Sheet (WAIS). It has been a tenet of the latter part of the twentieth century that the WAIS, known as a marine ice sheet because much of its bed is below sea level, is potentially unstable and may undergo disintegration if the fringing ice shelves were to disappear. However, the theoretical foundations of this belief are shaky and more careful consideration of the relevant
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66˚S
60˚W
65˚S
Scale 20
Kilometers 0 66˚S
62˚W
Figure 5 Modeled strain-rate trajectories on Larsen Ice Shelf superimposed on a Radarsat image taken on 21 March 1997. The trajectory of the smallest principal strain-rate (e_2 ) is negative (compressing flow, blue) over nearly all the Larsen Ice Shelf in its posticeberg calving configuration (January 1995) but is positive (red) in the region where the iceberg calved. The pattern shows isotropic points (‘stars,’ green) near where glaciers enter the ice shelf and, for the ‘iceberg’ area a ‘monstar’ (yellow cross) near the ice front. ‘Stars’ are seen for all the different ice front configurations, indicating that no fundamental change is occurring near the grounding line due to retreat of the ice front. A ‘monstar’ is only seen for Larsen B before the iceberg calved in 1995. The disappearance of the isotropic point is perhaps a ‘weak’ indication that the iceberg calving event was greater than expected for normal calving from a stable ice shelf and suggests further retreat may be expected.
dynamics suggests that the WAIS is no more vulnerable than any other part of the ice sheet. The problem lies in how to model the transition zone between ice sheet and ice shelf. If the transition is sharp, where the basal traction varies over lengths of the order of the ice thickness, then it can be considered as a passive boundary layer and does not affect the mechanics of the sheet or shelf to first order. Mass conservation must be respected but there is no need to impose further constraints. Requiring the ice sheet to have the same thickness as the ice shelf at the grounding line permits only two stable grounding line positions for a bed geometry which deepens inland. Depending on the depth of the bed
below sea level, the ice sheet will either shrink in size until it disappears (or, if the bed shallows again, the grounding line is fixed on a bed near sea level), or grow to the edge of the continental shelf. The lack of intermediate stable grounding line positions in this kind of model has been used to support the idea of instability of marine ice sheets. However, permitting a jump in ice thickness at the transition zone means the ice sheet system can be in neutral equilibrium, with an infinite number of steady-state profiles. Another kind of transition zone is a smooth one, where the basal traction varies gradually, for example, along an ice stream. The equilibrium dynamics have not been worked out for a full three-dimensional flow,
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but it seems that the presence of ice streams destroys the neutral equilibrium and helps to stabilize marine ice sheets. This emphasizes the importance of understanding the dynamics of ice streams and their role in the marine ice sheet system. Attempts to include moving grounding lines in whole ice sheet models suffer from incomplete specification of the problem. Assumptions built into the models predispose the results to be either too stable or too unstable. Thus there are no reliable models that can analyze the glaciological history of the Antarctic Ice Sheet. Predictive models rely on linearizations to provide acceptable accuracy for the near future but become progressively less accurate the longer the timescale.
See also Bottom Water Formation. Icebergs. Ice–ocean interaction. Sub Ice-Shelf Circulation and Processes.
Further Reading Doake CSM and Vaughan DG (1991) Rapid disintegration of Wordie Ice Shelf in response to atmospheric warming. Nature 350: 328--330.
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Doake CSM, Corr HFJ, Rott H, Skvarca P, and Young N (1998) Breakup and conditions for stability of the northern Larsen Ice Shelf, Antarctica. Nature 391: 778--780. Hindmarsh RCA (1993) Qualitative dynamics of marine ice sheets. In: Peltier WR (ed.) Ice in the Climate System, NATO ASI Series, vol. 12, pp. 67--99. Berlin: Springer-Verlag. Kellogg TB and Kellogg DE (1987) Recent glacial history and rapid ice stream retreat in the Amundsen Sea. Journal of Geophysical Research 92: 8859--8864. Robin G and de Q (1979) Formation, flow and disintegration of ice shelves. Journal of Glaciology 24(90): 259--271. Scambos TA, Hulbe C, Fahnestock M, and Bohlander J (2000) The link between climate warming and break-up of ice shelves in the Antarctic Peninsula. Journal of Glaciology 46(154): 516--530. van der Veen CJ (ed.) (1997) Calving Glaciers: Report of a Workshop 28 February–2 March 1997. BPRC Report No. 15. Columbus, Ohio: Byrd Polar Research Center, The Ohio State University. van der Veen CJ (1999) Fundamentals of Glacier Dynamics. Rotterdam: A.A. Balkema. Vaughan DG and Doake CSM (1996) Recent atmospheric warming and retreat of ice shelves on the Antarctic Peninsula. Nature 379: 328--331.
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IGNEOUS PROVINCES M. F. Coffins, University of Texas at Austin, Austin, TX, USA O. Eldholm, University of Oslo, Oslo, Norway
dynamic, nonsteady-state circulation within the Earth’s mantle, and suggests a strong potential for LIP emplacements to contribute to, if not instigate, major environmental changes.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1290–1298, & 2001, Elsevier Ltd.
Introduction Large igneous provinces (LIPs) are massive crustal emplacements of predominantly Fe- and Mg-rich (mafic) rock that form by processes other than normal seafloor spreading. LIP rocks are readily distinguishable from the products of the two other major types of magmatism, midocean ridge and arc, on the Earth’s surface on the basis of petrologic, geochemical, geochronologic, geophysical, and physical volcanological data. LIPs occur on both the continents and oceans, and include continental flood basalts, volcanic passive margins, oceanic plateaus, submarine ridges, seamounts, and ocean basin flood basalts (Figure 1, Table 1). LIPs and their small-scale analogs, hot spots, are commonly attributed to decompression melting of hot, low density mantle material known as mantle plumes. This type of magmatism currently representB10% of the mass and energy flux from the Earth’s deep interior to its crust. The flux may have been higher in the past, but is episodic over geological time, in contrast to the relatively steady-state activity at seafloor spreading centers. Such episodicity reveals
Composition, Physical Volcanology, Crustal Structure, and Mantle Roots LIPs are defined by the characteristics of their dominantly Fe- and Mg-rich (mafic) extrusive rocks; these most typically consist of subhorizontal, subaerial basalt flows. Individual flows can extend for hundreds of kilometers, be 10s to 100s of meters thick, and have volumes as great as 104–105 km3. Sirich rocks also occur as lavas and intrusive rocks, and are mostly associated with the initial and late stages of LIP magmatic activity. Relative to midocean ridge basalts, LIPs include higher MgO lavas, basalts with more diverse major element compositions, rocks with more common fractionated components, both alkalic and tholeiitic differentiates, basalts with predominantly flat light rare earth element patterns, and lavas erupted in both subaerial and submarine settings. As the extrusive component of LIPs is the most accessible for study, nearly all of our knowledge of LIPs is derived from lavas forming the uppermost 10% of LIP crust. The extrusive layer may exceed 10 km in thickness. On the basis of geophysical, predominantly seismic data from LIPs, and from comparisons with normal oceanic crust, LIP crust beneath the extrusive layer is believed to consist of
Figure 1 Phanerozoic global LIP distribution (red), with LIPs labeled (Table 1).
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IGNEOUS PROVINCES
an intrusive layer and a lower crustal body, characterized by compressional wave velocities of 7.0– 7.6 km s1, at the base of the crust (Figure 2). Beneath continental crust this body may be considered as a magmatically underplated layer. Seismic wave velocities suggest an intrusive layer that is most likely gabbroic, and a lower crust that is ultramafic. If the LIP forms on pre-existing continental or oceanic crust or along a divergent plate boundary, dikes and sills are probably common in the middle and upper crust. The maximum crustal thickness, including extrusive, intrusive, and the lower crustal body, of an oceanic LIP is B35 km, determined from seismic and gravity studies of the Ontong Java Plateau (Figure 1, Table 1). Low-velocity zones have been observed recently in the mantle beneath the oceanic Ontong Java Plateau, Oceanic plateau off-axis
x LCB
on-axis
x MC LCB
Ocean basin flood basalt x
LCB Continental flood basalt x Continental crust LCB Volcanic margin Rift Zone Postopening sediment COB x MC LCB
Continental crust
LIP CRUST Oceanic crust
Intrusives
Figure 2 Schematic LIP plate tectonic settings and gross crustal structure. LIP crustal components are: extrusive cover (X), middle crust (MC), and lower crustal body (LCB), continentocean boundary (COB), Normal oceanic crust is gray.
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as well as under the continental Deccan Traps and Parana´ flood basalts (Figure 1, Table 1). Interpreted as lithospheric roots or keels, the zones can extend to at least 500–600 km into the mantle. In contrast to high-velocity roots beneath most continental areas, and the absence of lithospheric keels in most oceanic areas, the low-velocity zones beneath LIPs apparently reflect residual chemical and perhaps thermal effects of mantle plume activity. High-buoyancy roots extending well into the mantle beneath oceanic LIPs would suggest a significant role in continental growth via accretion of oceanic LIPs to the edges of continents.
Distribution, Tectonic Setting, and Types LIPs occur worldwide, in both continental and oceanic crust in purely intraplate settings, and along present and former plate boundaries (Figure 1, Table 1), although the tectonic setting of formation is unknown for many features. If a LIP forms at a plate boundary, the entire crustal section is LIP crust (Figure 2). Conversely, if one forms in an intraplate setting, the pre-existing crust must be intruded and sandwiched by LIP magmas, albeit to an extent not resolvable by current geological or geophysical techniques. Continental flood basalts, the most intensively studied LIPs due to their exposure, are erupted from fissures on continental crust (Figure 1, Table 1). Most continental flood basalts overlie sedimentary basins that formed via extension, but it is not clear what happened first, the magmatism or the extension. Volcanic passive margins form by excessive magmatism during continental breakup along the trailing, rifted edges of continents. In the deep ocean basins, four types of LIPs are found. Oceanic plateaus, commonly isolated from major continents, are broad, typically flat-topped features generally lying 2000 m or more above the surrounding seafloor. They can form at triple junctions (e.g., Shatsky Rise), midocean ridges (e.g., Iceland), or in intraplate settings (e.g., northern Kerguelen Plateau). Submarine ridges are elongated, steep-sided elevations of the seafloor. Some form along transform plate boundaries, e.g., Ninetyeast Ridge. In the oceanic realm, oceanic plateaus and submarine ridges are the most enigmatic with respect to the tectonic setting in which they are formed. Seamounts, closely related to submarine ridges, are local elevations of the seafloor; they may be discrete, form a linear or random grouping, or be connected along their bases and aligned along a ridge or rise. They commonly form in
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Table 1
Large igneous provinces
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intraplate regions, e.g., Hawaii. Ocean basin flood basalts, the least studied type of LIP, are extensive submarine flows and sills lying above and postdating normal oceanic crust.
Ages Age control for all LIPs except continental flood basalts is sparse due to their relative inaccessibility, but the 40Ar/39Ar dating technique is having a particularly strong impact on studies of LIP volcanism. Geochronological studies of continental floor basalts (e.g., Siberian, Karoo/Ferrar, Deccan, Columbia River; Figure 1) suggest that most LIPs result from mantle plumes which initially transfer huge volumes (B105–107 km3) of mafic rock into localized regions of the crust over short intervals (B105–106 years), but which subsequently transfer mass at a far lesser rate, albeit over significantly longer intervals (107– 108 years). Transient magmatism during LIP formation is commonly attributed to mantle plume ‘heads’ reaching the crust following transit through all or part of the Earth’s mantle, whereas persistent magmatism is considered to result from steady-state mantle plume ‘tails’ penetrating the lithosphere which is moving relative to the plume (Figure 3). However, not all LIPs have obvious connections to mantle plumes or hot spots, suggesting that more than one source model may be required to explain all LIPs. LIPs are not distributed uniformly in time. During the past 150 million years for example, many LIPs formed between 50 and 150 million years ago, whereas few have formed during the past 50 million
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Large Igneous Provinces and Mantle Dynamics The formation of various sizes of LIPs in a variety of tectonic settings on both continental and oceanic lithosphere suggests a variety of thermal anomalies in the mantle that give rise to LIPs as well as strong lithospheric control on their formation. Equivalent mantle plumes beneath continental and oceanic lithosphere should produce more magmatism in the latter scenario, as oceanic lithosphere is thinner, allowing more decompression melting. Similarly, equivalent mantle plumes beneath an intraplate region (e.g., Hawaii) and a divergent plate boundary (e.g., Iceland) (Figure 1, Table 1) will produce more magmatism in the latter setting, again because decompression melting is enhanced. Recent seismic tomographic images of mantle plumes beneath Iceland and Hawaii show significant differences between the two. Only recently, seismic tomography has revealed that slabs of subducting lithosphere can penetrate the entire Earth’s mantle to the D00 layer at the boundary Crust Upper mantle
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years (Figure 4). Such episodicity likely reflects variations in rates of mantle circulation, and this is supported by high rates of seafloor spreading during a portion of the 50–150 million year interval. Thus, although LIPs manifest types of mantle processes distinct from those resulting in seafloor spreading, waxing and waning rates of overall mantle circulation probably affect both sets of processes. A major question that emerges from the global LIP production rate is whether the mantle is circulating less vigorously as the Earth ages.
660 km
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Figure 3 Model of the Earth’s interior showing plumes, subducting slabs, and two mantle layers that move in complex patterns, but never mix.
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km3 year
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Figure 4 LIP production, corrected for subduction and averaged over a 15 million year running window, since 150 million years ago (Ma).
between the mantle and core at B2900 km depth (Figure 3). If we assume that the volume of the Earth’s mantle remains roughly constant through geological time, then the mass of crustal material fluxing into the mantle must be balanced by an equivalent mass of material fluxing from the mantle to the crust. Most, if not all of the magmatism associated with the plate tectonic processes of seafloor spreading and subduction is believed, on the basis of geochemistry and seismic tomography, to be derived from the upper mantle (above B660 km depth). It is most reasonable to assume that the lithospheric material that enters the lower mantle is eventually recycled, in some part contributing to plume magmas.
Large Igneous Provinces and the Environment The formation of LIPs has had documented environmental effects both locally and regionally. The global
effects are less well understood, but the formation of some LIPs may have affected the global environment, particularly when conditions were at or near a threshold state. Eruption of enormous volumes of basaltic magma during LIP formation releases volatiles such as CO2, S, Cl, and F (Figure 5). A key factor affecting the magnitude of volatile release is whether eruptions are subaerial or submarine; hydrostatic pressure inhibits vesiculation and degassing of relatively soluble volatile components (H2O, S, Cl, F) during deep-water submarine eruptions, although low solubility components (CO2, noble gases) are mostly degassed even at abyssal depths. Investigations of volcanic passive margins and oceanic plateaus have demonstrated widespread and voluminous subaerial basaltic eruptions. Another important factor in the environmental impact of LIP volcanism is the latitude at which the LIP forms. In most basaltic eruptions, released volatiles remain in the troposphere. However, at high latitudes, the tropopause is relatively low, allowing
Increased planetary albedo
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large mass flux, basaltic fissure eruption plumes to transport SO2 and other volatiles into the stratosphere. Sulfuric acid aerosol particles that form in the stratosphere after such eruptions have a longer residence time and greater global dispersal than if the SO2 remains in the troposphere; therefore they have greater effects on climate and atmospheric chemistry. The large volume of subaerial basaltic volcanism, over relatively brief geological intervals, at highlatitude LIPs would contribute to potential global environmental effects. Highly explosive felsic eruptions, such as those documented from volcanic passive margins, an oceanic plateau (Kerguelen) (Figure 1, Table 1), and continental flood basalt provinces, can also inject both particulate material and volatiles (SO2, CO2) directly into the stratosphere. The total volume of felsic volcanic rocks in LIPs is poorly constrained, but they may account for a small, but not negligible fraction of the volcanic deposits in LIPs. Significant
volumes of explosive felsic volcanism would further contribute to the effects of plume volcanism on the global environment. Between B145 and B50 million years ago, the global oceans were characterized by variations in chemistry, relatively high temperatures, high relative sea level, episodic deposition of black shales, high production of hydrocarbons, mass extinctions of marine organisms, and radiations of marine flora and fauna (Figure 6). Temporal correlations between the intense pulses of igneous activity associated with LIP formation and environmental changes suggest a causal relationship. Perhaps the most dramatic example is the eruption of the Siberian flood basalts (Figure 1, Table 1) B250 million years ago, coinciding with the largest extinction of plants and animals in the geological record. Around 90% of all species became extinct at that time. On Iceland, the 1783–84 eruption of Laki provides the only human record of experience with the type of volcanism that
Figure 6 Temporal correlations among geomagnetic polarity, crustal production rates, LIPs, seawater strontium (Sr), sea level, climate, black shales, and extinctions.
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IGNEOUS PROVINCES
constructs LIPs. Although Laki produced a basaltic lava flow representing B1% of the volume of a typical (103 km3) LIP flow, the eruption’s environmental impact resulted in the deaths of 75% of Iceland’s livestock and 25% of its population from starvation.
Conclusions Oceanic plateaus, volcanic passive margins, submarine ridges, seamounts, ocean basin flood basalts, and continental flood basalts share geological and geophysical characteristics indicating an origin distinct from igneous rocks formed at midocean ridges and arcs. These characteristics include: (1) broad areal extent (>104 km2) of Fe- and Mg-rich lavas; (2) massive transient basaltic volcanism occurring over 105–106 years; (3) persistent basaltic volcanism from the same source lasting 107–108 years; (4) lower crustal bodies characterized by compressional wave velocities of 7.0–7.6 km s1; (5) some component of more Si-rich volcanic rocks; (6) higher MgO lavas, basalts with more diverse major element compositions, rocks with more common fractionated components, both alkalic and tholeiitic differentiates, and basalts with predominantly flat light rare earth element patterns, all relative to midocean ridge basalts; (7) thick (10s–100s of meters) individual basalt flows; (8) long (r750 km) single basalt flows; and (9) lavas erupted in both subaerial and submarine settings. There is strong evidence that LIPs both manifest a fundamental mode of mantle circulation commonly distinct from that which characterizes plate tectonics, and contribute episodically, at times catastrophically, to global environmental change. Nevertheless, it is important to bear in mind that we have literally only scratched the surface of oceanic, as well as continental LIPs.
See also Deep-Sea Drilling Results. Geophysical Heat Flow. Gravity. Magnetics. Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Seismic Structure. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Propagating Rifts and Microplates. Seamounts and Off-Ridge Volcanism. Seismic Structure
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Further Reading Carlson RW (1991) Physical and chemical evidence on the cause and source characteristics of flood basalt volcanism. Australian Journal of Earth Sciences 38: 525--544. Campbell IH and Griffiths RW (1990) Implications of mantle plume structure for the evolution of flood basalts. Earth and Planetary Science Letters 99: 79--93. Coffin MF and Eldholm O (1994) Large igneous provinces: crustal structure, dimensions, and external consequences. Reviews of Geophysics 32: 1--36. Cox KG (1980) A model for flood basalt vulcanism. Journal of Petrology 21: 629--650. Davies GF (2000) Dynamic Earth: Plates, Plumes and Mantle Convection. New York: Cambridge University Press. Duncan RA and Richards MA (1991) Hotspots, mantle plumes, flood basalts, and true polar wander. Reviews of Geophysics 29: 31--50. Hinsz K (1981) A hypothesis on terrestrial catastrophes: wedges of very thick oceanward dipping layers beneath passive continental margins – their origin and paleoenvironmental significance. Geologisches Jahrbuch E22: 3--28. Macdougall JD (ed.) (1989) Continental Flood Basalts. Dordrecht: Kluwer Academic Publishers. Mahoney JJ and Coffin MF (eds.) (1997) Large Igneous Provinces: Continental, Oceanic, and Planetary Flood Volcanism. Washington: American Geophysical Union Geophysical Monograph 100. Morgan (1981) Hotspot tracks and the opening of the Atlantic and Indian oceans. In: Emiliani C (ed.) The Oceanic Lithosphere, The Sea vol. 7, pp. 443--487. New York: John Wiley. Richards MA, Duncan RA, and Courtillot VE (1989) Flood basalts and hot-spot tracks: plume heads and tails. Science 246: 103--107. Saunders AD, Tarney J, Kerr AC, and Kent RW (1996) The formation and fate of large oceanic igneous provinces. Lithos 37: 81--95. Sigurdsson H, Houghton BF, McNutt SR, Rymer H and Stix J (eds.) (2000) Encyclopedia of Volcanoes, San Diego Academic Press: p. 1147. Sleep NH (1992) Hotspot volcanism and mantle plumes. Annual Review of Earth and Planetary Science 20: 19--43. White R and McKenzie D (1989) Magmatism at rift zones: the generation of volcanic continental margins and flood basalts. Journal of Geophysical Research 94: 7685--7729. White RS and McKenzie D (1995) Mantle plumes and flood basalts. Journal of Geophysical Research 100: 17543--17585.
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INDIAN OCEAN EQUATORIAL CURRENTS M. Fieux, Universite´ Pierre et Marie Curie, Paris Cedex, France Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1298–1308, & 2001, Elsevier Ltd.
Introduction Dynamically the equatorial area is a singular region on the earth because the Coriolis force is small, vanishing exactly at the equator. This results in a current structure that differs from that at other latitudes. Moreover, the equatorial current system of the Indian Ocean is entirely different from the current system found near the equator in the Pacific and Atlantic. This is principally due to its different wind forcing, which is described in the first section below. The systems of strictly equatorial currents at surface and at depth are reviewed in the second section. The third and fourth sections describe the North-East and South-West Monsoon Currents, north of the Equator, and the South Equatorial Countercurrent and the South Equatorial Current, south of the Equator.
component of the winds corresponding to the reversals between the NE and the SW monsoons, particularly in the western region along the Somali coast where the winds are the strongest. Between the monsoons, during the two transition periods, at the Equator, moderate eastward winds blow in spring (April–May) and in fall (October– November), with maxima between 701E and 901E (Figure 2). They could blow into one or several eastward bursts with a large seasonal and interannual variability. During the NE monsoon, the mean zonal component of the wind is weakly westward west of 801E, increasing in strength near the Somali coast. During the SW monsoon, the zonal components of the winds is weakly westward between 601E and 701E; west and east of that region, they are weakly eastward 40oE
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The winds over the tropical Indian Ocean are quite different from the winds over the Atlantic and the Pacific tropical oceans, where the NE and SE trade winds blow always in the same direction. Instead, the Indian Ocean (Figure 1), north of 101S, is under the influence of a monsoonal circulation, with complete reversal of the winds twice a year. The winter monsoon (December–March) blows from the NE in the Northern Hemisphere and from the NW south of the Equator toward the intertropical convergence zone (ITCZ) located near 101S. The change in direction at the Equator comes from the change of sign of the Coriolis force. The summer monsoon (June–September) blows from the SW in the Northern Hemisphere in continuity with the SE trade winds of the Southern Hemisphere, particularly in the western part of the equatorial ocean. The winds are stronger during the summer monsoon season. At the equator, during the monsoons, the winds have a preponderant meridional component, southward during the winter monsoon and northward during the summer monsoon, particularly near the western boundary. The result is a strong annual cycle in the meridional
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Figure 1 Winds over the Indian Ocean during the NE monsoon (A) and during the SW monsoon (B) with locations used in the text.
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Figure 2 Longitude time plots of equatorial zonal wind stress averaged from 11S to 11N, determined from the FSU pseudostress climatology for the period 1970–1996: (A) the total wind in dyn cm2; (B) time-averaged mean zonal wind stress; (C) annual zonal wind stress component; (D) semiannual zonal wind stress component. Contour interval is 0.05 dyn cm2, eastward wind stress regions are shaded. (From W Han, JP McCreary, DLT Anderson, AJ Mariano (1999) Dynamics of the eastern surface jets in the equatorial Indian Ocean. Journal of Physical Oceanography 29: 2191–2209, (1999) copyright by the American Meteorological Society.)
with a maximum between 801E and 901E. As a result, the annual mean zonal wind is eastward, the opposite of what is found in the other equatorial oceans, and is maximum around 801E to 851E. This particular wind regime is dominated by an annual and a semiannual cycle with similar amplitudes for the zonal components. The amplitude of the annual component is maximum near the western boundary associated with the stronger monsoon winds off Somalia. Relative annual component maxima are also found near 821E and near the eastern boundary due to the southward monsoon winds extension south of Sri Lanka and along the Indonesian coast. In contrast, the semiannual component of the equatorial zonal winds has a simple structure with a single maximum in the central ocean between 601E and 801E during spring and fall.
Currents at the Equator Currents in the Extreme West, along the Somali Coast
Compared to the other major oceans, there are very few direct current measurements in the Indian Ocean and most of the surface currents information comes
from the ship drifts and recently from satellitetracked drifting buoys. Along the Somali Coast, as seen above, the annual period is the dominant variability in the wind forcing. It is only relatively recently (1984–1986) that long-term direct current measurements were carried out at the equator within 200 km off the Somali coast. Above 150 m they show a seasonally reversing flow in phase with the NE and the SW monsoons, but a striking asymmetry between both monsoon seasons. While during the stronger SW summer monsoon the north-eastward Somali Current decays monotonically in the vertical, during the weaker NE winter monsoon the south-westward surface current is limited to the surface layer and a north-eastward countercurrent exists between about 150 m and 400 m, remnant of the SW monsoon current, followed again by weak south-westward current underneath down to about 1000 m (Figure 3). This sliced structure is confined to the equatorial region corresponding to the equatorial waveguide. Surface Currents at the Equator, East of 521E
Ship-drift climatology indicates that, at the equator, the surface currents reverse direction four times a
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INDIAN OCEAN EQUATORIAL CURRENTS
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Figure 3 Monthly mean current profiles observed on three moorings in the Somali current at the equator during the NE monsoon (January 1985), the SW monsoon (July 1985), and during the transition period (April 1985); monthly mean of component parallel to coast. The crosses at the surface are historical ship drift data from Cutler and Swallow (1984). The dotted near-surface profile in April is from the ship mounted acoustic Doppler current profiler. (From F Schott (1986). Seasonal variation of cross-equatorial flow in the Somali current. Journal of Geophysical Research 91(C9): 10581– 10584, copyright by the American Geophysical Union.)
year. During the two transition periods, under the influence of the equatorial eastward winds, a strong surface eastward jet (known as the ‘Wyrtki jet’ because Klaus Wyrtki first identified it by looking at different shipdrift atlases in the 1960s) sets up in a
narrow band, trapped in the equatorial wave guide within 2–31 of the equator (Figure 4). The Wyrtki jet exists mostly in the central and eastern regions and disappears in the western part where the currents have a strong meridional component as shown above. Owing to the efficiency with which zonal winds can accelerate zonal currents at the equator where the Coriolis force disappears, the current speed can reach 1 m s1. The ship drift climatology indicates roughly the same strength for the two jets, with the October–November one (1 m s1) slightly stronger than the April–May one (0.90 m s1). As there is a large seasonal and interannual variability in the eastward winds, strong currents could be found in April–May instead of October–November during some years. At the equator, in the middle of the Indian Ocean, the first long-term current measurements (1973–1975), near Gan Island (73110E–0141S), show energetic eastward currents throughout the upper
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Buoy 1885 20
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Figure 6 Trajectory of one drifting buoy (drogued at 20 m) entrained eastward into the equatorial Wyrtki jet in November–December 1981, then in the reversal in December–January, crossing the Arabian Sea, and undergoing the seasonal reversal of the current south of Sri Lanka. (From M. Fieux (1987) Oce´an Indien et Mousson ‘Unesco Anton Bruun Lecture’ 15 pages, unpublished manuscript.)
100 m mixed layer in phase with the zonal eastward component of local winds during the two transition periods between the monsoons (Figures 5, 6). The establishment of this eastward jet could be explained through eastward-propagating Kelvin waves triggered by the eastward-going winds during the transition periods. The stopping of the jet is thought to be the result of the reflection of the eastward-propagating Kelvin waves on the eastern coast into westward-propagating equatorially trapped Rossby waves that progressively impede the jet from east to west. This has been observed with drifting buoys (Figure 7). The eastward currents at the equator are convergent (contrary to the westward currents). Consequently, drifting buoys launched in the jet stay in it and are a good means of observing its variability. In Figure 8 the drifting buoys were entrained into the May jet. They then stopped during the SW monsoon drifting slowly south of the equator and were taken up again eastwards into the November jet. Between the strong eastward flow periods, the equatorial surface currents are westward and much weaker (Figures 4 and 6). The associated change of current direction during each transition period produces semiannual variations in the thermocline depth and sea level. During periods of eastward flow when warm water is carried toward the east, the thermocline deepens off Sumatra and rises off Africa, corresponding to opposite displacements of the sea surface. Strong eastward flow at the equator entrains a surface convergence and as a result a downwelling in the upper layer, contrary to what is found in the
Figure 7 Drifting buoy trajectories between 601E and Sumatra during the period of September–February. Full lines are for 1979– 1980 and dashed lines are for 1980–1981. Dashed arrows indicate different speeds on this time–longitude diagram. During the first year the reversal propagation speed toward the west was around 55 cm s1, and around 40 cm s1 during the second year. (From G Reverdin, M Fieux, J Gonella and J Luyten (1983) Free drifting buoy measurements in the Indian Ocean equatorial jet. In: Nihoul JCJ (ed.) Hydrodynamics of the Equatorial Ocean. Amsterdam: Elsevier Science, 99–120.)
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Figure 8 (A) Satellite-tracked drifting drogued (at 20 m) buoys trajectories in 1979; dots on the trajectories are spaced at 15-day intervals. (B) Wind vectors estimated from cloud trajectories on satellite images during the equatorial eastward wind period (24 May 1979) and during the fully developed SW monsoon (19 June 1979). (The short tick on the arrow corresponds to 5 knots, the longer one to 10 knots, and the triangle to 50 knots.) (Reprinted with permission from JR Luyten, M Fieux and J Gonella. Equatorial currents in the Western Indian Ocean, Science 209: 600–603, copyright 1980 American Association for the Advancement of Science.)
other oceans. It is only during the weak westward flow that there is a weak upwelling. Long-term direct current measurements carried between Sri Lanka and 01450 S, in 1993–1994, reveal a large seasonal asymmetry that year in the semiannual eastward jet transport, with 35 Sv in November 1993 (surface velocities exceed 1.3 m s1) and only 5 Sv in May 1994. These transports were calculated with a lower boundary sets at 200 m as, sometimes, the eastward currents are not confined to one core but extend into the ray-like structures of the equatorial waves that continue to greater depths as in November 1993 (Figure 9). This large variability is due to a large seasonal and interannual variability of the zonal winds partly related to the Southern Oscillation. For example, during the El Nin˜o of the century in 1997, the equatorial eastward winds in the Indian Ocean completely disappeared in October–November and were replaced by westward winds. Numerical model results show that direct wind forcing is the dominant forcing mechanism of the equatorial surface jets. The semiannual response of the current to the wind is nearly three times as strong as the annual one, despite similar amplitudes in the corresponding wind components. This is due to the simpler zonal structure of the semiannual wind and to the resonance of the basin to the semiannual component. The mixed layer shear flow seems also to
enhance the semiannual response. The jet is strengthened when the mixed layer is reduced, particularly when the precipitation during the northern hemisphere summer and fall thins the mixed layer in the eastern ocean and so could strengthen the November jet in the east. The reflection of equatorial Kelvin waves into westward going Rossby waves at the eastern boundary is also important in weakening and even canceling the directly forced eastward jet about 2 months after the wind onset. The presence of the Maldives islands blocks part of the equatorially trapped waves. The effect on the semiannual waves is to weaken both jets, and the effect on the annual waves tends to weaken the May jet and to strengthen the November jet. Currents at Depth at the Equator
It is only during the NE monsoon, the season when the large-scale wind structure resembles the Pacific and the Atlantic ones, that a similar eastward equatorial undercurrent (EUC) embedded in the thermocline exists. Observations of the equatorial undercurrent are scarce. It was first observed during the International Indian Ocean Expedition (IIOE) in March–April 1963, between 531E and 921E, with speeds up to 0.8 m s1, then at Gan Island (731E) in March 1973 with velocities of up to 1 m s1, but not the following year. In 1975 and 1976, it was found
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Figure 9 Monthly mean zonal velocities in cm s1 for the upper 300 m, south of Sri Lanka, along 801300 E, between August 1993 and August 1994. Contour interval is 10 cm s1 and shaded areas indicate eastward currents. Note the unusual reappearance of the EUC in August 1994. (From J Reppin, F Schott and J Fischer (1999). Equatorial currents and transports in the upper central Indian Ocean: Annual cycle and interannual variability, Journal of Geophysical Research 104(C7): 15495–15514, copyright by the American Geophysical Union.)
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INDIAN OCEAN EQUATORIAL CURRENTS
from January extending into May–June at 551300 E with observed speeds reaching 0.8 m s1 in February and March at the end of the NE monsoon (Figure 10). Direct measurements give a maximum transport of 17 Sv, in March–April 1994, at the longitude of Sri Lanka (801300 E) 0 (Figure 9). It is confined between 2130 N–21300 S with slight meandering and is weak east of 801E. Its core is around 100 m in the upper equatorial thermocline. During the NE monsoon, it flows under a weak westward current until April–May when the eastward Wyrtki jet starts, then the whole upper layer flows eastward. It stops in May–June but surprisingly, in 1994, it reappeared in August. The existence of the EUC is related to equatorial westward winds, which force a westward surface current that builds up a zonal pressure gradient below the mixed layer that maintains the eastward undercurrent. The slope of the sea surface is opposite to the slope of the thermocline. The reappearance of the undercurrent in August 1994 is effectively related to anomalous onset of westward winds in the eastern part of the ocean, during May and June 1994. They force a westward surface current, again building a subsurface zonal pressure gradient, and thus an eastward undercurrent reappeared with some delay in August 1994. In 1976, current profiles show that the vertical velocity structure in the vicinity of the equator is characterized by small vertical scales of order of 50– 100 m in the upper layer increasing with depth throughout the water column. This deep jetlike structure is trapped to within 11 of the Equator and has a timescale of the order of several months. One-year current-meter measurements made at 200 m, 500 m, and 750 m at the Equator, between 481E and 621E, and recently at 801E, present a dominant semiannual reversal of the zonal component at all depths, much deeper than could be explained by direct wind forcing (Figure 11). Furthermore, these reversals do not happen at the same time at different depths. This seasonal cycle, with larger vertical scales, has been shown to penetrate vertically. The measurements suggest a mixture of equatorial Kelvin waves and long equatorial Rossby waves propagating downward from the surface where they are forced by the winds. The zonal velocity shows upward phase propagation, and downward energy propagation away from the surface. The behavior of the zonal currents is characteristic of an eastward-propagating equatorial Kelvin wave and a westward-propagating long equatorial Rossby wave. The ratio of the semiannual energy in the east current component to that in the north component is 40 to 1. This shows how
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Figure 10 Meridional section along 531E of zonal velocities in February, March, and April 1975 showing the equatorial undercurrent. Contours interval is 20 cm s1. Plain contours correspond to eastward currents and dashed contours to westward currents. (From A Leetmaa and H Stommel (1980) Equatorial current observations in the Western Indian Ocean in 1975 and 1976, Journal of Physical Oceanography, 10: 258–269, copyright by the American Meteorological Society.)
much the equatorial ocean is a barrier to the meridional motions.
Currents North of the Equator South of Sri Lanka
North of the Equator, the monsoon forcing drives a general eastward flow, the South-west Monsoon Current (SMC), during the fully developed SW monsoon, and a general westward flow, the Northeast Monsoon Current (NMC), during the NE monsoon, extending south to about 21S in January– February. The exchanges between the Arabian Sea and the Bay of Bengal are restricted to the south of India and Sri Lanka. Drifting-buoy trajectories show
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Figure 11 (A) Vertical section of the Indian Ocean along the Equator with the location of the current-meter array between 471E and 621E. (B) Time series of east velocity from seven current-meters at 750 m depth starting on April 1 1979. (C) Contour plot of lowfrequency east velocity as a function of depth and time (day 01 April 1979). At each nominal depth (200, 500, 750 m), four records are filtered by a 30-day running mean and averaged together. The contour interval is 5 cm s1. (From JR Luyten and DH Roemmich (1982) Equatorial currents at semiannual period in the Indian Ocean, Journal of Physical Oceanography 12: 406–413, copyright by the American Meteorological Society.)
the seasonal current reversal in that restricted region (Figure 6). Direct observations carried out in 1991–1994 south of Sri Lanka show that the monsoon currents are mostly confined in the upper 100 m. In August 1993, the eastward SMC extended to the equator and retracted north of 41300 N in September. It was replaced in October by the westward NMC, extending south to about 21N, with speeds between 0.3 and 0.8 m s1 (Figure 9). Shipdrifts show that, around the Maldives, the NMC splits into a branch that bends south-westward and a branch that follows the western coast of India, bringing low-salinity Bay of Bengal water into the Arabian Sea. The NMC maximum flow lies north of 41N with a mean transport of 10–12 Sv (Figure 12). In May the current reverses eastwards into the SMC again. The SMC transport reaches about 8–10 Sv with speeds up to 0.75 m s1 in July. The SMC is sometimes separated from the coast of Sri Lanka by a coastal westward countercurrent bringing Bay of Bengal water. The annual mean flow past Sri Lanka was 2–3 Sv westward in 1993–1994. Numerical models show
that these monsoon currents are driven by the largescale tropical wind field. Contrary to the equatorial circulation where the semiannual period prevails, the annual component dominates the upper layer flow north of 41N.
Currents South of the Equator Apart from shipdrifts and satellite-tracked drifting buoys, there are no direct long-term current measurements in the South Equatorial Countercurrent and only two in the South Equatorial Current. The South Equatorial Countercurrent
In contrast to the Pacific and the Atlantic oceans, where a countercurrent is found year-long north of the Equator between the North Equatorial Current and the South Equatorial Current, in the Indian Ocean the countercurrent is found south of the Equator and has a large extent only during a short period of the year.
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Figure 12 Seasonal variability of the transport in the upper 300 m from moorings south of Sri Lanka, between 31450 N and 51520 N, in 1991–1992. (The solid circles are from ship-mounted acoustic Doppler current profiler measurements in December 1990–January 1991; the solid circle in parentheses is for July 1991, and the cross is from Pegasus profiler measurements in March 1992.) (From F Schott, J Reppin, J Fisher and D Quadfasel, Currents and transports of the monsoon current south of Sri Lanka, Journal of Geophysical Research 99(C12): 25127–25141, copyright by the American Geophysical Union.)
During the winter monsoon, the Somali current flows southward, crosses the Equator, and merges with the northward flowing East Africa Coastal Current (EACC), at about 2–41S, to form the eastward South Equatorial Countercurrent (SECC). It is a region of high eddy activity. The SECC is found between 21S and 81S during January–March. During that season the winds have an eastward component at those latitudes just north of the ITCZ. The speeds vary between 0.5 and 0.8 m s1 in the west, getting weaker in the east. In March 1995, at 801E, observed surface velocities exceed 0.7 m s1 and the current extended to about 1100 m, with an eastward geostrophically deduced transport of about 55 Sv. In the east, part of it continues into the Java Current and part of it recirculates southward into the South Equatorial Current (SEC). It is detected in the meridional slope of the thermocline which slopes downward toward the Equator, with an opposite upward slope of the sea surface (Figure 13). Together with the SEC and the EACC, the SECC forms an elongated cyclonic gyre. In the latitude range about 10–131S and between 201S and 251S, recent measurements of the sea level variability through satellite altimetry show westward propagation of sea level anomalies corresponding to semiannual and annual Rossby wave characteristics. The wind-driven model results show westward propagation of Rossby waves in the shear zone between the SECC and the SEC that are obstructed and partially reflected by the Mascarenes banks
(55–601E). In the west, at the end of the winter season, in March–April, when the eastward winds start on the equator, the outflow from the EACC into the SECC begins to move northward toward the Equator and the eastward flow at that time is mostly equatorial.
The South Equatorial Current
The westward South Equatorial Current forced by the south-east trade winds extends south of the SECC. It represents the northern branch of the South Indian subtropical gyre. It is seen in the meridional downward slope of the thermocline towards the south (Figure 13). It is partly fed, in its northern part, between 101S and 141S, by the low-salinity throughflow jet originating from the western Pacific Ocean through the Indonesian Seas carrying about 4 to 12 Sv with larger extremes depending on the year. The low salinity extends down to about 1200 m. In March 1995, at 801E, measured surface velocities reached 0.7 m s1. The SEC is the limit of the influence of the monsoon system and separates the northern and southern Indian Ocean. It is stronger during July–August when the SE trade winds are stronger. Its indirectly computed mean transport relative to the 1000 m level varies from 39 Sv in July– August to 33 Sv in January–February with large uncertainty. Its mean transport increases from east to west. Its latitudinal range varies between 8–221S in
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INDIAN OCEAN EQUATORIAL CURRENTS
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Figure 13 Meridional temperature section along 54–551E in May 1981 with 11C isotherm interval. The downward slope of the thermocline form 91S toward the south corresponds to the SEC, and the downward slope toward the Equator corresponds to the SECC. Close to the Equator, the accentuated slope corresponds to the May Wyrtki jet. (From M Fieux and C Levy (1983) Seasonal observations in the western Indian Ocean. In: Nihoul JCJ (ed.) Hydrodynamics of the Equatorial Ocean. Amsterdam: Elsevier Science, 17–29.)
Buoy 1886 10
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Figure 14 Satellite-tracked drifting buoy trajectory between November 1981 and February 1984. The buoy, drouged at 20 m, was entrained eastward in the November jet then in the SECC and back westward into the SEC. One year later it was again entained in the same elongated gyre. (From M Fieux (1987) Oce´an Indien et Mousson, ‘Unesco Anton Bruun Lecture’, unpublished manuscript.)
July–August and 10–201S in January–February. Its velocities range between 0.3 m s1 and 0.7 m s1. The SEC impinges both on the east coast of Madagascar and on the east African coast, resulting
in several intensified boundary currents along these coasts. East of Madagascar, the SEC splits near 171S into a northward flow, carrying about 27 Sv near 121S, and a southward flow, transporting 20 Sv
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between the surface and 1100 m near 231S. At 121S energetic boundary current transport variations occur at the 40–55-day period, contributing to about 40% to the total transport variance, while at 231S the 40–55-day period fluctuations contribute only 15% to the total transport variance. The northern branch of the SEC splits again east of the African coast near Cape Delgado (111S) into the southward flowing Mozambique Current (MC) and the northward flowing East African Coastal Current (EACC). Drifting-buoy trajectories describe an elongated cyclonic gyre of the equatorial current system composed of the equatorial eastward jet or the SECC, depending on the season, the off Sumatra SE current, the westward SEC, and the EACC. Some drifting buoys, launched during a transition period at the Equator, carried into the eastward Wyrtki jet and in the SECC, crossed the basin, then were driven southeastward into the Sumatra current, then crossed the basin westward into the SEC, and flowed back to the Equator in the western region exactly one year later when they were again carried into the SECC (Figure 14).
is eastward, associated with a strong convergence at the surface inducing an equatorial downwelling. The upwelling regions are found instead north of the equator, along the Somalia, the Arabian, and the Indian coasts, and are seasonally depending. The equatorial current structure in the Indian ocean is complex and further long-term observations as well as modeling efforts are needed to better understand its seasonal and interannual variability and its role in the large-scale meridional exchanges between the northern and the southern Indian Ocean.
See also Agulhas Current. Coastal Trapped Waves. Current Systems in the Indian Ocean. Elemental Distribution: Overview. El Nin˜o Southern Oscillation (ENSO). Rossby Waves. Somali Current. Water Types and Water Masses. Wind Driven Circulation.
Further Reading Conclusion Owing to stronger winds in the tropical Indian ocean, the currents are stronger than in the Pacific and Atlantic Oceans but seasonally are highly variable. Away from the western boundary the equatorially trapped long waves (Kelvin and Rossby) explain most of the observed seasonal variations of the equatorial currents. In contrast to the Atlantic and Pacific Oceans, equatorial upwelling is weak in the Indian Ocean. During the period of strong current, the surface flow
Cutler AN and Swallow JC (1984) Surface Currents of the Indian Ocean (to 251S, 1001E): Compiled from Historical Data Archived by the Meteorological Office, Brack-nell, UK. Wormley, UK: Institute of Oceanographic Science (unpublished report) 187, 8pp. and 36 charts. Fein JS and Stephens PL (eds.) (1987) Monsoons. Washington DC: NSF and Wiley. Open University, Oceanography Course Team (1993) Ocean Circulation. Oxford: Pergamon Press. Tomczak M and Godfrey S (1994) Regional Oceanography: An Introduction. Oxford: Pergamon Elsevier.
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INDONESIAN THROUGHFLOW J. Sprintall, University of California San Diego, La Jolla, CA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The Indonesian seas play a unique role in providing the only open pathway that connects two major ocean basins at tropical latitudes. On average, the sea level is higher on the western Pacific side of the Indonesian archipelago compared to the eastern Indian side. This pressure gradient generates a transport of water and their properties from the Pacific toward the Indian Ocean. This flow from the Pacific Ocean into the Indian Ocean is known as the Indonesian Throughflow. The warmer water that enters from the Pacific can be traced throughout the Indonesian seas, and then followed within the surface to intermediate depths as a distinct low-salinity tongue in the Indian Ocean. As such, the Throughflow forms the ‘warm’ water route for the global thermohaline circulation and therefore impacts the regional and global climate system. Because of its proximity to Asia and Australia, the circulation and transport in the Indonesian seas has a large seasonal variation due to the influence of the reversing annual wind patterns associated with the Asian–Australian monsoon system. During the different monsoon seasons, waters of different sources from both the Indian and Pacific Oceans flow into the Indonesian archipelago and cause variability in temperature, salinity, and other properties. Local processes within the Indonesian seas related to the regional monsoon winds such as upwelling and downwelling, along with the tides, air–sea heat fluxes, and voluminous precipitation and associated river runoff, also act to change the Pacific temperature and salinity stratification into the distinctly fresh Indonesian seas profile. These changes in the physical properties of the water are linked to the behavior, migration pattern, and the seasonal distribution of the phytoplankton and pelagic fish species that live within the Indonesian seas. Thus a knowledge and understanding of the pathways and variability of the Indonesian Throughflow and its properties is important for the region’s people, who depend on the sea for their very food and livelihood, and also to help develop management plans to sustain these valuable and limited maritime resources.
Pathways and Water Masses The combination of numerous islands, with narrow straits that connect a series of large, deep ocean basins within the Indonesian seas, provides a winding route for the Indonesian Throughflow (Figure 1). The surface to upper thermocline waters in the Indonesian seas are primarily drawn from the North Pacific subtropical waters and North Pacific Intermediate Water that flows southward in the Mindanao Current, east of the Philippines. These waters mostly take the ‘western’ Throughflow route through the Sulawesi Sea (sometimes known as the Celebes Sea) into the Makassar Strait. Within Makassar Strait, the 650-m-deep Dewakang sill permits only the upper thermocline waters to enter the Flores and Banda Seas, or to directly exit into the Indian Ocean via the shallow (350 m) Lombok Strait. Smaller contributions of North Pacific surface water may also take the ‘eastern’ Throughflow route, through the Maluku Sea and over the deeper (1940 m) sill of Lifamatola Strait. Lower thermocline and deeper water masses – Antarctic Intermediate Water (AAIW) and Circumpolar Deep Water – that originate in the South Pacific in the New Guinea Coastal Undercurrent, also enter the Indonesian seas via the eastern route through the Lifamatola Strait. These deeper water masses are saltier than the upper waters that originate from the North Pacific. Upper waters of South Pacific origin can also flow directly over the Halmahera Sea (blocked below 700-m depth) and into the internal Seram and Banda Seas. Most of the transformation of the Pacific water masses into the distinctive Indonesian water profile of relatively isohaline water from the thermocline to near bottom occurs within the Banda Sea. As the different water masses flow toward the Banda Sea over the various Indonesian sills and straits, the abrupt changes in bottom topography cause intense diapycnal mixing. The region-average vertical diffusivities at the sills are greater than 10 3 m2 s 1, which is an order of magnitude larger than both that typically estimated for the interior ocean, and also for that which occurs within the interior of the Banda Sea itself. At least some of this diapycnal mixing is attributable to the strong tides within the Indonesian seas. At the sills, mixing can also occur via entrainment and shear-induced turbulence, including internal wave generation. It is also likely that hydraulic effects can restrict the flow and affect the circulation pattern in some straits due to the shoaling
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Figure 1 Map of the Indonesian region showing the bathymetry and pathways of the mean Indonesian Throughflow. Shallow and upper thermocline pathways from the North Pacific are shown by the red arrows, and lower thermocline to intermediate depth pathways are shown by the yellow arrows. Islands, seas, and straits referred to in the text are indicated. The 200-m isobath is shown by the dashed line.
and contraction of sills, although as yet there have been no suitable direct observations or process studies undertaken that may resolve this response. Density-driven (or isopycnal) flow dominates the deeper layers of South Pacific origin that enter through Lifamatola Strait into the Banda Sea, driving deep circulation patterns that keep the internal basins well ventilated. In the surface to upper thermocline layer, the diapycnal mixing combines with the tidal forcing, wind-driven upwelling, air–sea heat flux, and surface precipitation and river runoff to form the low-salinity Indonesian Throughflow Water (ITW: 34–34.5 psu). Below this sits the Indonesian Intermediate Water (IIW), which has a silica maximum and is slightly saltier (36 psu) as a result of diapycnal mixing of the intermediate Pacific water masses with a larger contribution from the saltier water masses of the South Pacific found at depth in the Banda basin. From the Banda Sea, the Indonesian Throughflow exits into the southeast Indian Ocean via the Timor Passage (1250 m eastern sill depth at Leti Strait, and 1890 m at the western end), or through Ombai Strait (sill depth 3250 m) and then through the Savu Sea to
Sumba Strait (900 m) and Savu Strait (1150 m). The Indonesian Throughflow waters are apparent as a freshwater jet in the westward-flowing South Equatorial Current (SEC) across the entire tropical Indian Ocean between 81 and 171 S. In the Indian Ocean, the total Throughflow is recognizable as the lowsalinity ITW in the upper 100–200 m of surface water and the deeper-salinity minimum of the IIW which is found at B600–1200-m depth, and hence above the eastern sill depth of Timor Passage. Although the ITW and IIW are separate cores even directly west of Timor Passage as they exit the Indonesian seas, they gradually increase in salinity and become even more distinct entities as the highsalinity maximum subtropical and equatorial waters of the open Indian Ocean encroach between them. The ITW and IIW lose their low-salinity signature more slowly than the intervening salinity maximum during westward advection across the Indian Ocean because the adjacent water masses on their isopycnals (surface layer and AAIW) are considerably fresher than other Indian Ocean water masses. Nonetheless, the slight increase in salt results in an increase in density and thus the Throughflow waters
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become deeper and erode toward the western Indian Ocean, although the latitude of both the core ITW and IIW remains remarkably constant within the SEC because of the nearly exact zonality of potential vorticity in the Tropics. At the western edge of the Indian Ocean, it is likely that much of the Throughflow enters the Agulhas Current system, although because of mixing with other water masses in this region and beyond, it becomes more difficult to directly trace their origin from the Indonesian seas. Nonetheless, numerous Agulhas eddies in the Atlantic have been identified as containing relatively fresh Indonesian thermocline water, providing observational evidence to support the numerical models that show the Indonesian Throughflow playing an important role in the global circulation.
Mean Throughflow Mass and Heat Flux Estimates The magnitude and variability of the Indonesian Throughflow are still sources of major uncertainty for both the modeling and observational oceanography communities. They are the dominant sources of error in the basin-wide heat and freshwater budgets for the Pacific and Indian Oceans. Though general circulation models are gradually improving, they are unable at present to reproduce the narrow passages and convoluted bottom topography of the internal Indonesian seas in order to adequately resolve the structure and variability of the ITF transport. Thus the models still largely rely on the scantly available observations to provide guidance in their initialization and validation for simulations of the ocean circulation and climate. Historical estimates of the mean mass transport of the Indonesian Throughflow are wide-ranging, from near zero to 30 Sv (1 Sv ¼ 106 m3 s 1). In part, this sizeable difference is due to the lack of direct measurements, but also because of the real variation that can severely alias mean estimates if survey periods are not sufficiently long. Recent mooring measurements of velocity in Makassar Strait in 1986–87 suggest a volume transport of around 8 Sv for the surface to upper thermocline waters that pass through this inflow channel. Long-term direct measurements of the lower thermocline to intermediate water masses have yet to be obtained, although water mass considerations and synoptic survey measurements of the Throughflow at the eastern edge of the Indian Ocean suggest a volume transport of 2–3 Sv. A 3-month current meter record at Lifamatola Strait in early 1985 determined a 1.5-Sv contribution of deep water to the Throughflow. Direct long-term measurements
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of the Throughflow within the exit passages have also been obtained at different times over the past few decades. The transport through Timor Passage (1988–89) and Ombai Strait (1995) appears to be roughly equally partitioned, at c. 4–6 Sv through each strait, whereas the transport through the shallower exit passage of Lombok Strait (1985) is B2 Sv. Hence the total volume transport associated with the fulldepth Throughflow is probably around 10–14 Sv, although because of the different sampling years of the different direct measurement programs this total volume transport estimate should be treated with caution. As will be discussed below, the volume and properties of the Throughflow are known to change with various phases of the El Nin˜o–Southern Oscillation (ENSO) cycle, and hence this may have impact on the transport estimates from a given passage in a given year. Only with multiyear, simultaneous measurements of the full-depth velocity structure in all the major Throughflow passages can the volume transport be accurately determined. Direct measurements that meet these criteria to determine transport will become available from an array of moorings that were deployed as part of the International Nusantara Stratification and Transport (INSTANT) program. The 11 moorings were deployed in the major inflow and outflow passages in December 2003, turned around in June 2005, and are to be recovered in December 2006. The heat and fresh water carried by the Indonesian Throughflow potentially impact the basin budgets of both the Pacific and Indian Oceans. Earlier estimates of the heat transport ranged from 0.5 to 1.0 PW (petawatt, with 1 PW ¼ 1015 W) as the Throughflow was assumed to be surface-intensified, and reflect the relatively warm temperatures found in its western Pacific source waters. However, the recent observations from the Makassar moorings suggest that the strongest Throughflow velocities (southward in this strait) are found in the thermocline, and hence the cool transport-weighted temperature of B15 1C leads to a mean heat transport of B0.4–0.5 pW (relative to 0 1C). Including an as yet unresolved colder component from the IIW would only further reduce the Throughflow impact on the heat transport of the Southern Hemisphere. Because of the lack of subsurface salinity measurements, the contribution of the freshwater flux from the Indonesian Throughflow to the Indian Ocean and beyond is even less well constrained. However, since the salinity of the Pacific waters entering the Indonesian seas is fresher than the inflow from the Southern Ocean northward into the Pacific, the Throughflow probably represents a net transport of fresh water out of the Pacific and into the Indian Ocean where the
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salinity increases before being exported southward in the Agulhas Current.
Variability in the Throughflow Properties and Transport Pacific, Indian, and regional Indonesian wind forcing modulates the Indonesian Throughflow mass and property transport over a wide range of timescales that combine to produce variability of the same order of magnitude as the long-term mean estimates cited above. In this section, we will discuss the cause of the different timescales of variability and their impact on the Throughflow. Because it is the main driving force of the Throughflow we begin with a discussion of the seasonal variability related to the Asian–Australasian monsoon regime. Forcing at other timescales from remote Pacific and Indian Ocean sources generally tends to modulate this fundamental seasonal response. Monsoonal-related Variability: Annual Timescales
The large-scale sea level difference between the easterly trade winds in the western Pacific and the reversing monsoonal winds over the southern Indonesian seas drives the annual change in the Throughflow. During the northwest monsoon from December to February, south of the equator the winds over Indonesia are to the southeast (Figure 2(a)). Through Ekman dynamics, the southeastward winds cause a drop in sea level along the northern Nusa Tenggara islands and warm surface waters accumulate in the Banda Sea, acting to reduce the Throughflow transport into the Indian Ocean. The downwelling south of Nusa Tenggara leads to warmer surface waters here. The convective rainfall associated with the northwest monsoon spans the Indonesian archipelago to produce heavy rainfall and subsequently high river runoff that together combine to freshen the surface waters, particularly in the Java Sea region. The flow from the Banda Sea toward the Indian Ocean is strongest during the southeast monsoon, from June through September, when the winds are to the northwest and more intense (Figure 2(b)). The strong winds during this monsoon result in a lower sea level to the south of Nusa Tenggara, and hence the flow through the exit passages is thought to be enhanced by this local Ekman response. Surface waters are cooler in the Banda Sea, and also south of Nusa Tenggara due to the winddriven upwelling. The convective activities move northward of the equator so conditions in Indonesia are drier and surface salinity higher compared to the
northwest monsoon. The wind-driven regime also leads to changes in the chlorophyll concentrations, which is elevated in regions of upwelling as nutrients from the deep are bought to the surface, and reduced where there is downwelling. Chlorophyll indicates phytoplankton abundance and therefore gives a broad view of the biological activity and associated fishery distribution during the different monsoon seasons. While the surface response to the reversing monsoonal winds is clearly evident in Figure 2, the monsoonal effect on the Throughflow velocity and transport is much more complicated. While all the observations presently available in the exit passages do indeed show the expected maximum Throughflow occurring during the southeast monsoon period, the Makassar Strait mooring data showed the maximum southward flow occurred during the 1996–97 northwest monsoon period. The large differences in peak transport timing between the inflow and outflow straits are most likely due to storage of waters in the Banda Sea on seasonal timescales. It appears that the Banda Sea acts as a reservoir, filling up and deepening the thermocline during the northwest monsoon. During the more intense southeast monsoon, Ekman divergence in the Banda Sea combined with the lower sea level off the south coast of the Nusa Tenggara are more conducive to drawing waters into the Indian Ocean. The maximum Ekman convergence (or gain) of 1.7 Sv causes the thermocline to heave almost 40 m during October–November, although there is matching divergence (or loss) occurring in April–May of each year. Convergences and divergences of this magnitude can have a substantial impact on thermocline stratification, and thus may affect the sea surface temperature by changing the temperature of the water being entrained into the mixed layer. In addition, since the Banda Sea is also the primary site for conversion of the Pacific waters into the distinct Indonesian stratification profile, it is likely that the composition and magnitude of the stored waters could have a significant impact on the Indian Ocean heat, fresh water, and mass budgets. The response of the surface properties to the shifting monsoonal winds has recently been suggested to play another important role in modifying the Throughflow velocity and transport. The buoyant low-salinity water present in the Java Sea and southern Makassar Strait during the northwest monsoon (Figure 2(a)) creates a northward pressure gradient that inhibits the warm surface water in Makassar Strait from flowing southward into the Indian Ocean, even though the predominant winds are southward during this monsoon. During the
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INDONESIAN THROUGHFLOW
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Figure 2 The average wind stress (N m 2) and sea surface height (cm; top panels), sea surface temperature (1C; middle panels) and sea surface salinity (psu; bottom panels) during the (a) northwest monsoon (Dec.–Feb.) and (b) southeast monsoon (Jun.–Aug.). The 200-m isobath is shown by the dotted lines in the lower panel.
southeast monsoon, the surface waters are more saline (Figure 2(b)), the northward pressure gradient is eliminated, and the northward winds tend to constrain the surface Throughflow. Thus, it appears that during both monsoon seasons, the strongest southward Throughflow occurs at thermocline depths
(100–200 m) in Makassar Strait, resulting in the recently determined cooler ocean heat transport cited in the above section. Although the data are still preliminary, the INSTANT moorings in the shallower Lombok Strait also show a subsurface southward velocity maximum at B50 m during the
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southeast monsoon. During the northwest monsoon in Lombok Strait the main southward flow is pushed even deeper to B100 m, as strong northward flows are found at the surface. Whether these northward surface flows are related to remote Indian Ocean or local wind forcing is further discussed in the following section. Variability Related to Indian Ocean Forcing: Intraseasonal and Semiannual Timescales
Remote wind forcing from the Indian Ocean is responsible for variability in the Indonesian Throughflow transport and properties via the generation of coastal Kelvin waves. Anomalous wind bursts in the equatorial Indian Ocean force an eastward equatorial jet associated with an equatorial Kelvin wave that, upon impinging the west coast of Sumatra, results in poleward propagating coastal Kelvin waves. Westerly wind bursts force downwelling Kelvin waves semiannually during the monsoon transitions, nominally in May and November, and intraseasonally (30–90 days) most likely in response to the intraseasonal wind forcing of the Madden–Julian Oscillation. The poleward propagation of the downwelling coastal Kelvin wave raises the sea level along the Indonesian wave guide of Sumatra and Nusa Tenggara, transporting warm and fresh surface water and causing maximum eastward flow of the boundary current that flows along southern Nusa Tenggara, the South Java Current. While there is some discrepancy between theoretical models as to whether the Indian Ocean-forced Kelvin wave energy can penetrate northward through Lombok Strait, the available observational data clearly show the northward flow of warm water that indicates the presence of Kelvin waves in both Lombok Strait and Makassar Strait, particularly during the May transition period. Furthermore, some of the Kelvin wave energy is also readily apparent in observations from Ombai Strait, further east along the coastal waveguide, during the May and November transition periods, suggesting that Indian Ocean water can penetrate through the Savu Sea and beyond. Commensurate with vertical ray-tracing theory, the semiannual energy from remotely forced Kelvin waves extends deeper in the water column with distance along the waveguide, and so the reversals in flow are evident well below the thermocline in all three straits. What proportion of the Kelvin wave energy is ‘lost’ through Lombok Strait versus that remaining in the coastal waveguide, and how this may be modulated by the annual cycle, is the focus of active research. Similarly, whether the northward surface flows observed during the
northwest monsoon in the INSTANT data from Lombok Strait and Ombai Strait, and to some extent in Makassar Strait, are related to Kelvin wave intrusions or to locally generated wind forcing is also the subject of ongoing research. The northward flows during this monsoon phase have a clear intraseasonal variability that may be related to the remote winds associated with the intraseasonal Madden–Julian Oscillation in the eastern equatorial Indian Ocean. However, given the relatively shallow vertical penetration of these reversals, they could also be a locally wind-driven Ekman dynamical response.
Variability Related to Pacific Ocean Forcing: Interannual Timescales
Remote winds in the Pacific Ocean drive large-scale circulation that impacts the Indonesian Throughflow on low-frequency, interannual timescales. The weakening of Pacific trade winds that occurs during El Nin˜o reduces the sea level gradient between the Pacific and Indian Oceans that is thought to be the driving force for the Throughflow. The changes are of the sense that mass and heat transport are smaller and the thermocline is shallower during El Nin˜o, with the converse being true during La Nin˜a time periods. Within the inflow Makassar Strait, the 1996–98 mooring velocity and temperature time series show a strong correlation with the Southern Oscillation Index that tracks the ENSO-related variability in the Pacific. Along the Nusa Tenggara exit passages, the relationship of transport and temperature to ENSO variability is much less certain, primarily because we lack the multiyear, full-depth measurements that are required to establish such a relationship. Surface geostrophic transport variability (0–100-m depth) over the period 1996–99 through the exit passages, inferred from the cross-strait pressure gradient measured by shallow pressure gauges, suggest diminished flow through Lombok and Ombai Straits during the El Nin˜o years 1997–98, with increased flow through Timor Passage. This is consistent with modeling studies, in which there is stronger flow through Timor Passage and weaker flow through Ombai Strait during El Nin˜o events. Interestingly, the surface geostrophic transport through Lombok Strait was strongly northward during the northwest monsoon of the 1997–98 El Nin˜o event, amplified over the expected northward flows in this strait during this monsoon phase. It is possible, as discussed above, that the main southward Throughflow occurred at depth during this time period and hence was not measured by the shallow pressure gauges.
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INDONESIAN THROUGHFLOW
Interannual variability in sea level is also evident within the Banda Sea and along the Northwest Australian shelf region down to around 201 S. Both models and observations suggest that the remote Pacific tropical winds associated with ENSO drive both equatorial and off-equatorial Rossby waves that are responsible for the observed sea level and associated thermocline response. The Rossby wave signal is radiated through the Banda Sea, as well as entering the coastal waveguide at the western tip of New Guinea, and propagates poleward through Timor Strait and along the Australian shelf as coastally trapped waves. In turn, these trapped waves excite free westward propagating Rossby waves at the Australian coast that can be detected several hundreds of kilometers offshore in the southeast tropical Indian Ocean. Interannual Kelvin waves forced remotely by equatorial interannual zonal winds in the Indian Ocean have also been documented as impacting the thermocline and sea level variability along the Nusa Tenggara coastal waveguide, similar to those observed on intraseasonal and semiannual timescales. Again, both models and observations suggest these Kelvin waves penetrate through Lombok Strait and into the internal Indonesian seas. Thus, as suggested by Wijffels and Meyers, the Indonesian seas comprise the intersection of two oceanic wave guides from remotely generated wind forcing in both the Pacific and Indian Oceans on interannual timescales.
Conclusions The mass and property characteristics of the Indonesian Throughflow vary over the full range of tidal to interannual timescales. Within the internal Indonesian seas, the Pacific temperature and salinity stratification, as well as the local sea surface temperature, are modified by the strong air–sea heat and freshwater fluxes, seasonal wind-induced upwelling, and large tidal forces. These ocean circulation and processes in the Indonesian seas influence not only the local climate but also the global climate through connections with the Pacific and Indian Ocean. For example, ruinous drought conditions in Australia have been linked to cool sea surface temperature anomalies in the
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Indonesian region. Similarly basin-wide anomalies in sea surface temperature, wind, and precipitation that spanned the Indonesian region during 1997–98 resulted in severe drought in Indonesia and catastrophic floods in eastern Africa. Thus, undoubtedly the oceanic heat and freshwater flux of the Indonesian Throughflow into the Indian Ocean affect the atmospheric– ocean coupling and are strongly linked to the evolution of interannual climate anomalies such as occur through the ENSO and monsoon systems. Understanding the variation in the Indonesian Throughflow is therefore crucial for understanding the coupled air– sea climate system, and the storage of the heat and fresh water that are ultimately redistributed throughout the world oceans by thermohaline circulation.
See also Coastal Trapped Waves. Current Systems in the Indian Ocean. Ekman Transport and Pumping. Heat and Momentum Fluxes at the Sea Surface. Ocean Circulation: Meridional Overturning Circulation. Rossby Waves. Upper Ocean Heat and Freshwater Budgets. Water Types and Water Masses. Wind Driven Circulation.
Further Reading Gordon AL (2001) Interocean exchange. In: Seidler G, Church J, and Gould J (eds.) Ocean Circulation and Climate, ch. 4.7, pp. 303--314. New York: Academic Press. Talley LD and Sprintall J (2005) Deep expression of the Indonesian Throughflow: Indonesian intermediate water in the South Equatorial Current. Journal of Geophysical Research 110: C10009 (doi:10.1029/2004JC002826). The Oceanography Society. (2005). Special Issue: The Indonesian Seas. Oceanography 18(4), 144pp. Wijffels S and Meyers GA (2004) An intersection of oceanic wave guides: Variability in the Indonesian Through flow region. Journal of Physical Oceanography 34: 1232--1253. Wyrtki K (1961) Physical oceanography of the Southeast Asian waters. Scripps Institution of Oceanography NAGA Report 2. San Diego, CA: Scripps Institution of Oceanography.
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE T. D. Dickey, University of California, Santa Barbara, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1313–1323, & 2001, Elsevier Ltd.
Introduction Light is of great importance for the physics, chemistry, and biology of the oceans. In this article, a brief introduction is provided to the two subdisciplines focusing on light in the ocean: ocean optics and biooptics. A few of the problems addressed by these subdisciplines are described. Several of the in situ sensors and systems used for observing the subsurface light field and optical properties are introduced along with general explanations of the operating principles for measuring optical variability in the ocean. Some of the more commonly used ocean platforms and optical systems are also discussed. Finally, some examples of oceanographic optical data sets are illustrated. Solar radiation, which includes visible radiation or light, impinges on the surface of the ocean. On average, a small fraction or percentage is reflected back into the atmosphere (roughly 6% on average) while a high fraction penetrates into the ocean. This fraction (or percentage), defined as the albedo, varies in time and space as a function of several factors including solar elevation, wave state, surface roughness, foam, and whitecaps. Radiative transfer is a branch of oceanography termed ‘ocean optics,’ a term that denotes studies of light and its propagation through the ocean medium. Radiative transfer processes depend on the optical properties of the components lying between the radiant source (e.g., the sun) and the radiation sink (the ocean and its constituents). Another commonly used term is ‘bio-optics,’ which invokes the notion of biological effects on optical properties and light propagation and vice versa. Some of the data used for examples here focus on bio-optical and physical interactions. Solar radiation spans a broad range of the electromagnetic energy spectrum. The visible portion, roughly 400–700 nm, is of primary concern for ocean optics for many practical problems. These include: underwater visibility, photosynthesis and primary production of phytoplankton, upper ocean
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ecology, red tides, pollution, biogeochemistry including carbon cycling, photochemistry, light-keyed migration of organisms, heating of the oceans, global climate change, and remote sensing of ocean color. It is worth noting that optical oceanographers are also interested in the ultraviolet (UV) range and the infrared. For example, UV radiation can damage phytoplankton and can also lead to modifications in genetic material, and thus evolution, whereas the penetration of the infrared is critical for near surface radiant heating. Solar radiation is clearly a primary driver for ocean physics, chemistry, and biology as well as their complex interactions. Thus, it is not surprising that virtually all important oceanographic problems require interdisciplinary approaches and necessarily atmospheric, physical, chemical, biological, optical, and geological data sets. These data sets should be collected concurrently (concept of synopticity) and span sufficient time and space scales to observe the relevant processes of interest. For local studies involving optics, variability at timescales of one day and one year are especially important. For global problems, variability extends well over ten orders of magnitude in space [O(millimeters) to O(104 kilometers)] and much longer in time for climate problems. Multiplatform observing approaches are essential. Capabilities for obtaining atmospheric and physical oceanographic data are relatively well advanced in contrast to those for chemical, biological, optical, acoustical, and geological data. This is not surprising, because of the greater complexity and nonconservative nature of the chemistry and biology of the oceans. Yet, remarkable advances are being made in these areas as well. In fact, several bio-optical, chemical, geological, and acoustical variables can now be measured on the same time and space scales as physical variables; however many more variables still need to be measured.
Fundamentals of Ocean Optics A few operation definitions need to be introduced before discussing the various optical sensors. Light entering the ocean can be absorbed or scattered. The details of these processes comprise much of the study of ocean optics. First, it is convenient to classify bulk optical properties of the ocean as either inherent or apparent. Inherent optical properties (IOPs) depend
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE
only on the medium and are independent of the ambient light field and its geometrical distribution. Inherent optical properties include: spectral absorption coefficient, aðlÞ, spectral scattering coefficient, bðlÞ, and spectral beam attenuation coefficient (or, sometimes denoted as ‘beam c’), cðl). To define these coefficients, consider a beam of monochromatic light impinging perpendicularly on a thin layer of water. The fraction of the incident radiant flux, F0 (energy or quanta per unit time), which is absorbed by the medium, Fa/F0, divided by the thickness of the layer is the absorption coefficient, a, in units of m1. Similarly, the analogous scattering coefficient, b (m1) is the fraction of the incident radiant flux scattered from its original path (primarily forward), Fb/F0, divided by the thickness of the layer. The beam attenuation coefficient is simply c ¼ a þ b. The spectral volume scattering function, bðc; l) represents the scattered intensity of light per unit incident irradiance per unit volume of water at some angle c (with respect to the exiting, nonscattered incident beam) into solid angle element DO. Units for bðc; l) are m1 sr1. This is essentially the differential scattering cross-section per unit volume in the parlance of nuclear physics. The spectral scattering coefficient is obtained by integrating the volume scattering function over all directions (solid angles). The forward scattering coefficient, bf , is obtained by integrating over the forward-looking hemisphere (c ¼ 0 to p/2) and the backward scattering coefficient, bb, is calculated by integrating the backlooking hemisphere (c ¼ p=2 to p). The spectral volume scattering phase function is the ratio of bðc; l) to bðlÞ with units of sr1. Other important IOPs include spectral single-scattering albedo, o0 ðlÞ ¼ bðlÞ=cðl), and fluorescence, which can be considered a special case of scattering. (Note: fluorescence is not strictly an IOP; however it is often used as a proxy for chlorophylla.) The proportion of light which is scattered versus absorbed is characterized by o0 ðl); that is, if scattering prevails then o0 ðl) approaches a value of 1 and if absorption dominates, o0 ðl) approaches 0. IOPs obey simple mathematical operations. For many applications, it is often convenient to partition the total absorption coefficient, at ðl), the total scattering coefficient, bt ðl), and the total beam attenuation coefficient, ct ðlÞ, in terms of contributing constituents such that: at ðlÞ ¼ aw ðlÞ þ aph ðlÞ þ ad ðlÞ þ ag ðlÞ
½1
bt ðlÞ ¼ bw ðlÞ þ bph ðlÞ þ bd ðlÞ þ bg ðlÞ
½2
ct ðlÞ ¼ cw ðlÞ þ cph ðlÞ þ cd ðlÞ þ cg ðlÞ
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½3
The subscripts indicate contributions by pure sea water (w), phytoplankton (ph), detritus (d), and gelbstoff (g) with units of m1 for all variables. Detritus is the term for particulate organic debris (e.g., fecal material, plant and animal fragments, etc.) and gelbstoff (sometimes called yellow matter or gilvin) is the term for optically active dissolved organic material. Note that eqns [1]–[3] are intended for use where bottom and coastal sediment contributions are minimal. The spectral absorption coefficient of pure sea water is well characterized and is nearly constant in space and time with greater absorption in the red than blue portions of the visible spectrum. The magnitudes of the detrital and gelbstoff absorption spectra tend to decrease monotonically with increasing wavelength and can be modeled. Phytoplankton spectral absorption varies significantly in relation to community composition and environmental changes; however characteristic peaks are typically found near wavelengths of 440 nm and 683 nm. Much bio-optics research focuses on the temporal and spatial variability of aph ðl). Ship-based ocean water samples have often been used for studies of the IOPs, however, it is quite preferable to obtain in situ measurements to insure representative local values as well as to characterize temporal and spatial variability, preferably with simultaneous physical, chemical, and biological measurements.
Instrumentation Beam transmissometers have been the most commonly used instruments for measuring IOPs with a variety of applications ranging from determinations of suspended sediment volume to phytoplankton biomass and productivity to particulate organic carbon (POC). The principle of operation involves the measurement of the proportion of an emitted beam, C, which is lost through both absorption and scattering as it passes to a detector through some pathlength DL. The beam attenuation coefficient, c, is then given as ½lnð1 CÞ=DL; or light intensity, I, is given as IðDLÞ ¼ Ið0Þ expðcDL). Similar expressions apply for absorption coefficient, a, and scattering coefficient, b. Beam transmissometers have often used a red light emitting diode (660 nm) for the collimated light source and pathlengths of 25 or 100 cm (long pathlengths are preferable for clearer waters). The wavelength of 660 nm was selected in order to minimize attenuation by gelbstoff, which attenuates strongly at shorter wavelengths but minimally in the red. One of the measurement
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complications for beam c is that most scattering in ocean waters is in the forward direction. The acceptance angle of the detector is thus made as small as possible (o 1–21) to minimize underestimation of beam c. Corrections are done for this effect as well as for instrument temperature and pressure effects. In situ multispectral attenuation–absorption (acmeters) instruments have recently become commercially available. These concurrently measure spectral absorption and attenuation coefficients, which are spectral signatures of both particulate and dissolved material. These devices use dual light sources and interference filters on a single rotating wheel (nine wavelengths from 400 to 715 nm with 10 nm bandwidth). Whereas earlier beam transmissometers were open to sea water, the newer ac-meters use two tubes through which sea water is pumped. The inside of the c-tube is flat black to minimize reflections whereas the a-tube is reflective (shiny) in order to maximize internal reflection to better estimate absorption. The deployment of these meters requires minimal lateral and torsional stressing in mounting (sensitivity of beam alignment) and good flow through the tubes. Calibrations involve temperature, salinity, and pure water measurements. A correction is also needed to account for the fact that not all scattered light is collected (causing a biased overestimate of absorption). Similar, though more capable instruments, have recently been developed to measure a and c at 100 wavelengths from 400 to 726 nm wavelengths with 3.3 nm resolution. A single white light source is used with fiberoptics to provide light to each of the two tubes (for a and b) as well as for a reference path (for correcting changes in lamp output). Each light path has its own spectrometer (using a 256 pixel photodiode array). Biofouling is an issue for all optical instruments. Closed (pumped) systems are less susceptible to fouling as they are primarily in the dark. A variety of special antifouling methods have been attempted with mixed success. These have included copper screens, copper flow tubes, bromide solutions, and other cleaning agents. Spectral scattering coefficient can be computed from ac-meter data by simply performing the difference b ¼ c a. Spectral a c meter data and relevant spectral decomposition models can be used to provide in situ estimates of aph ðl) and other components (e.g., gelbstoff) for a broad range of environmental conditions as a function of time at fixed depths using moorings or as a function of depth from ships or profilers deployed from moorings. The volume scattering function is one of the more important and challenging measurements of optical oceanography. Until recently, few instruments existed for this measurement and data collected with an
instrument developed in the 1970s were commonly used for many problems. The challenge arises because dominant scattering typically occurs at small angles (roughly 50% in the forward direction between 0 and 2–61) making it difficult to sense the weak scattering light, which is near the intense illuminating beam. Ideally, measurements should be done at several angles to better resolve the function. However, the backscatter signal, bb, can be estimated with a few measurement angles provided the form of the volume scattering function is relatively well known. Recently developed volume scattering and backscatter instruments take advantage of this function. An important consideration is that backscattering is related to the size distribution, shapes, and composition of the particles sampled (e.g., coastal versus open ocean particles). Thus, this type of information can be obtained in principle. Radiative transfer models need both absorption and volume scattering function information to estimate the propagation of light. Also, an important remote sensing parameter, remote sensing reflectance (discussed below) is related to bb and a. Key wavelengths are selected for the spectral measurements on the basis of the application and water types (e.g., coastal versus open ocean). Particle size distributions can also be measured using laser (Frauenhofer) diffraction instruments. These devices employ charged coupled device (CCD) array photodetectors (linear or circular ring geometries). Modified versions of some of these instruments can also measure particle settling velocities. Particles in the range of 5–500 mm can be measured with resolution dependent on the number of individual detector rings. These devices can also measure beam transmission by sensing the undeviated light. Irradiance is the radiant flux per unit surface area (units of W m2 or quanta (or photons) s1 m2 or mol quanta (or photons) s1 m2. Note that 1 mol of photons is 6.02 1023 (Avogadro’s number) and that one mole of photons is commonly called an einstein. It is convenient to define downwelling irradiance, Ed, as the irradiance of a downwelling light stream impinging on the top face of a horizontal surface (e.g., ideally flat light collector oriented perpendicular to the local gravity vector) and upwelling irradiance, Eu, as the irradiance of an upwelling light stream impinging on the bottom face of a horizontal surface. Irradiance reflectance, R, is Eu/Ed. Irradiance measurements are fundamentally important for quantifying the amount of light available for photosynthesis and for radiative transfer theory and computations. Another useful radiometric variable is radiance, L ¼ Lðy; FÞ, which is defined as the radiant flux at a specified point in a given direction per unit solid
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE
angle per unit area perpendicular to the direction of light propagation (units of W (or quanta s1) m2 sr1). Zenith angle, y, is the angle between a vertical line perpendicular to a flat plane and an incident light beam, and azimuthal angle, F, is the angle with respect to a reference line in the plane of the flat plate. By integrating Lðy; FÞcosy over the solid angle ðdo ¼ siny dy dFÞ, of the upper hemisphere (say of a flat plate from y ¼ 0 to p/2 and F from 0 to 2p), one obtains Ed. Similarly, upwelling irradiance, Eu, is calculated by integrating over the bottom hemisphere. Net downward irradiance is the difference Ed Eu, or the integral of Lðy; FÞcosy over the entire sphere (full solid angle 4p). Scalar irradiance, E0, is defined as the integral of Lðy; FÞ over the entire sphere. It should be noted that all definitions are for a specific wavelength of light. If one integrates scalar irradiance over the visible wavelengths (roughly 400–700 nm; note that the visible range is sometimes defined as 350–700 nm), then the biologically important quantity called photosynthetically available radiation (PAR) is obtained. Apparent optical properties (AOPs) depend on both the IOPs and the angular distribution of solar radiation (i.e., the geometry of the subsurface ambient light field). To a reasonable approximation, the attenuation of spectral downwelling incident solar irradiance, Ed ðl; zÞ, can be described as an exponential function of depth: Ed ðl; zÞ ¼ Ed ðl; 0 Þexp½Kd ðl; zÞz
½4
where z is the vertical coordinate (positive downward), Ed ðl; 0 Þ is the value of Ed just below the air– sea interface, and Kd ðl; zÞ is the spectral diffuse attenuation coefficient of downwelling irradiance. Kd ðl) (here depth dependence notation is suppressed for convenience) is one of the important AOPs. The contributions of the various constituents (e.g., water, phytoplankton, detritus, and gelbstoff) to Kd ðl) are often represented in analogy to the absorption coefficients given in eqn [1][1]. This leads to the term ‘quasi-inherent’ optical property as it is suggestive that inherent optical properties, i.e., like at ðl), are closely related to Kd ðlÞ, which is often described as a quasi-inherent optical property. However, a key point is that Kd ðlÞ is in fact dependent on the ambient light field. Analogous spectral attenuation coefficients are defined for diffuse upwelling irradiance, downwelling radiance, and upwelling radiance, KL ðlÞ. It should be noted that for periods of low sun angle and/or when highly reflective organisms or their products (e.g., coccolithophores and coccoliths) are present, then multiple scattering becomes increasingly more important and simple relations
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between IOPs and AOPs break down. It is important to mention that spectral radiometric measurements of AOPs have been far more commonly made than measurements of IOPs. The relationships between IOPs and AOPs are central to developing quantitative models of spectral irradiance in the ocean. Radiative transfer theory is used to provide a mathematical formalism to link IOPs and the conditions of the water environment and light forcing to the AOPs of the water column. One motivation of this research is the desire to predict AOPs given environmental forcing conditions and estimates or actual measurements of IOPs. Another related, inverse goal is to be able to estimate or determine the vertical (and ideally horizontal) structure of IOPs and their temporal variability given normalized water-leaving radiance, Lwn ðlÞ, measured remotely from satellite or airplane color imagers. The extrapolation of subsurface values of LðlÞ to the surface has been the subject of considerable research, because a major requirement of ocean color remote sensing is to match in situ determinations of those from satellite- or plane-based spectral radiometers which necessarily must account for the effects of clouds and aerosols. As mentioned above, the important remote sensing parameter, spectral radiance reflectance, Lu ðlÞ=Ed ðl), has been found to be proportional bb/(a þ bb) or bb/a from Monte Carlo calculations for waters characterized by b/a values ranging from 1.0 to 5.0. The spectral radiance reflectance evaluated just above the ocean surface is defined as the remote sensing reflectance Rrs ðl). It should also be noted that remote sensing estimates of pigment (typically chlorophyll a or chlorophyll a þ phaeopigments) concentrations use empirical relations involving ratios such as Lwn(443 nm)/ Lwn(550 nm) or Rrs(490 nm)/Rrs(555 nm) or these ratios with other wavelength combinations. One of the most commonly used instruments for measuring AOPs is a broadband scalar irradiance E0 or PAR sensor, because of the need for determining the availability of light for phytoplankton and its relative simplicity. A spherical light collector made of diffusing plastic receives the light from approximately 4p sr. The light is transmitted via a fiberoptic connector or quartz light conducting rod to a photodetector which records an output voltage. The instrument is calibrated with a standard lamp. Analogous sensors use flat plate cosine or hemispherical collectors. Spectroradiometers are irradiance or scalar irradiance meters which use a variable monochromator placed between a light collector and the photodetector. Light separation can be achieved using sets of interference filters (usually 10 nm in bandwidth) selected for particular purposes (e.g., absorption peaks and hinge points for pigments).
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Alternatively, higher spectral resolution (B3 nm from 400 to 800 nm) can be achieved using grating monochromators. Cosine (flat-plate), spherical, or hemispherical collectors are used for several different measurements or calculations of the AOPs described above (e.g., Ed, Lu, Kd, etc.). Irradiance cosine collectors are designed such that their responses to parallel radiant flux should be proportional to the angle between the normal to the collector surface and the direction of the radiant flux. This is necessary as the angle between the incoming radiant flux and the collector is variable, as is the area of the projection. To account for the differences in refractive indices of air and water, an immersion correction factor (typically 1.3–1.4) is applied. It should be noted that radiance sensors are designed to accept light in a small solid angle (typically a viewing angle of a few degrees). Irradiance and radiance instruments are calibrated using well-characterized, standard light sources. Because of the growing concern about UV radiation, a number of instruments are designed to measure into the UV portion of the electromagnetic spectrum as well. A grand challenge of oceanography is to greatly increase the variety and quantity of ocean measurements. Optical and other ocean measurements are expensive. Thus, a major goal is to develop new sensors and systems, which can be efficiently deployed from a host of available ocean platforms including ships, moorings, drifters, floats, and autonomous underwater vehicles (AUVs). A further need is to telemeter the data in near real-time. Shipbased observations are useful for detailed profiling (high vertical resolution), but moorings are better suited for high temporal resolution, long-term measurements. Drifters, floats, and AUVs can provide horizontal coverage unattainable from the other in situ platforms. Ultimately, all of these platforms, along with satellite- and plane-based systems, are needed to fill in the time–space continuum. Several optical systems, which have been used for mooringbased time series studies, are illustrated in Figure 1. These collective instrument systems include most of the sensors described earlier. The sampling for these instruments is typically done every few minutes (in some cases once per hour) for periods of several months. At this point in time, biofouling, rather than power or data storage, is the limiting factor. However, even this aspect is becoming less problematic with several new antibiofouling methods. In addition, more optical sensors are being deployed from other emerging platforms such as AUVs. The optimal utilization of optical data from the various platforms will require the use of advanced data merging methods and data assimilation models for predictions.
Optical Experiment Data Sets Finally, a few examples of data, which have been collected with some of the optical instruments described above (Figure 1), are presented. An interdisciplinary experiment devoted to the understanding of relations between mixing processes and optical variability was conducted south of Cape Cod, Massachusetts on the continental shelf. Several spectral optical instruments collected IOP and AOP data sets using ship-based profiling and towed systems, moorings, and bottom tripods. The period of the experiment covered almost one year (July 1996– June 1997). Some instruments sampled at several minute to hourly intervals whereas others sampled with vertical resolution on the scale of centimeters for a few weeks during two intensive field campaigns. Several interesting observations were enabled. These included: sediment resuspension forced by two passing hurricanes (hurricane Edouard and hurricane Hortense), spring and fall phytoplankton blooms, water mass intrusions, and internal solitary waves. Time series data collected using the BIOP system (e.g. BIOPS and MORS in Figure 1) which was deployed from a mooring and a bottom tripod located near each other in 70 m waters, are shown in Figure 2. Figure 2(A) shows the 37-m time series of total spectral absorption (water absorption has been subtracted) using a nine-wavelength ac-meter. The major feature is related to hurricane Edouard. The spectral absorption contributions due to phytoplankton are shown for two days in the summer in Figure 2(B); the peaks at 440 and 683 nm are caused by phytoplankton. Figure 2(C, D) shows ac-meter time series of spectral (nine wavelengths) scattering and attenuation coefficients. The record is dominated by sediment resuspension caused by mixing and waves created by hurricane Edouard and the more distant hurricane Hortense. The complete mooring data set showed that sediments were lifted more than 30 m above the ocean floor. The second example highlights optical data collected from a ship-based profiling system (similar to the system shown in Figure 1A). The purpose of the measurements was to study the dispersion of contaminants (treated wastewaters) discharged about 4 km offshore (70 m water depth) into Mamala Bay, off the coast of Honolulu, Hawaii in the fall of 1994. Optical and physical measurements were made from shipboard using both profile and towing modes in the vicinity of the outfall plume. Optical instrumentation included a beam transmissometer (660 nm), spectral ac-meters (nine wavelengths for a and c), PAR sensor, chlorophyll fluorometer, a particle size analyzer (laser diffraction method), and a
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE
spectral absorption and fluorescence instrument. Physical measurements of temperature, salinity and pressure were also done. Sampling was designed to track both the horizontal and vertical structure of effluent as manifest in the optical signals. The water contributions to a, b, and c were removed for the analyses. The collective optical and physical measurements enabled the partitioning of particle types into categories: particulate versus dissolved components, phytoplankton opposed to detrital components, shallow layer versus deep layer phytoplankton, and old versus newly discharged sewage plume waters. Briefly, a profile (see Figure 3) taken near the end of the outfall
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diffuser (water depth B 80 m) displayed the following features: (1) sewage plume waters centered near 60 m as characterized by low salinity and very high values of spectral attenuation and absorption (greater in the shorter wavelengths), and (2) a shallow phytoplankton layer near 20 m with modest relative maxima in chlorophyll fluorescence and spectral absorption and attenuation coefficients in the blue. The profile of spectral single-scattering albedo, o0 ðlÞ ¼ bðlÞ=cðlÞ, shows a general trend of increased importance of scattering for greater wavelengths with the most pronounced effect for the sewage plume waters centered at 60 m (Figure 3C).
(A)
(B) Par sensor
PAR sensor Pump
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3 Ed's UCSB data logger/ Battery pack 3 Lu's Fluorometer & VSF sensors
MORS
MOSS
Figure 1 Schematic showing a variety of optical systems for mooring applications. (A) System used for measuring inherent optical properties (IOPs) including beam c (660 nm), spectral attenuation and absorption coefficients at nine wavelengths, along with PAR and temperature. (B) System used for measuring inherent optical properties including spectral attenuation and absorption coefficients at 100 wavelengths, spectral backscatter at six wavelengths, along with PAR, temperature, and pressure. (C) System measuring spectral fluorescence with 6 excitation wavelengths and 16 emission wavelengths; also PAR and temperature sensors are included. (D) System for measuring apparent optical properties (AOPs) including spectral downwelling irradiance and upwelling radiance at seven wavelengths and PAR along with temperature. A telemetry module is also included. (E) System for measuring apparent optical properties (AOPs) including spectral downwelling irradiance and upwelling radiance at three wavelengths along with instruments for measuring chlorophyll fluorescence, volume scattering function. An antifouling shutter system is also utilized.
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37 m
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Figure 2 Illustrations showing data collected from the BIOPS system shown in Figure 1(A) on the continental shelf south of Cape Cod Massachusetts in the summer and fall of 1996. (A) A time series of the total spectral absorption coefficient of light (after subtracting the clear water component) at a depth of 37 m (total water depth is 70 m). (B) The spectral absorption on YD 204 and 239. (C) Time series of spectral scattering coefficient computed by differencing the total spectral attenuation and absorption coefficients (again, the clear water coefficient has been subtracted) at 68 m. The large peaks are attributed to passages of hurricanes Edouard (E) and Hortense (H). (D) as for (C) except for spectral beam attenuation coefficients. (Figures based on Chang and Dickey, 1999.)
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE
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Figure 3 Data collected near an ocean sewage outfall in Mamala Bay off Honolulu, Hawaii in the fall of 1994. Vertical profiles of (A) spectral absorption coefficient, (B) spectral attenuation coefficient, and (C) single scattering albedo. Data were obtained from an acmeter system similar to the one shown in Figure 1(A). Eight wavelengths are displayed and the major peak near 60 m is caused by sewage plume waters with high scattering and absorption coefficents. (Figures based on Petrenko et al., 1997.)
The final example describes an important linkage between in situ and remote sensing observations of ocean color. The Bermuda Testbed Mooring (BTM) is used to test new oceanographic instrumentation (including systems shown in Figure 1), for scientific studies devoted to biogeochemical cycling and climate change, and for groundtruthing (verification using in situ data) and algorithm development for ocean color satellites such as SeaWiFS. The latter aspect is the focus here. Two BTM optical systems utilizing radiometers for measurements of spectral downwelling irradiance and spectral upwelling radiance are illustrated in Figure 1(D) and (E). These systems measure at wavelengths, which are coincident with those of the SeaWiFS ocean color satellite, and sample at hourly intervals. Two or more systems are deployed at different depths and another radiometer system is deployed from the surface buoy to collect incident
spectral downwelling irradiance. Derived quantities include time series of several of the AOP quantities introduced earlier (e.g., spectral diffuse attenuation coefficients, spectral reflectances including remote sensing reflectance, and water-leaving radiance). Time series of spectral upwelled radiance from the BTM radiometer system (Figure 1D) located at 14 m are shown for the period April through July, 1999 in the top set of panels of Figure 4. The variability is primarily caused by the daily cycle of solar insolation, cloud cover, and optical properties associated with phytoplankton and their products. The bottom panels of Figure 4 show a subset of the water-leaving radiance data as determined from the BTM radiometers (extrapolation to surface) and SeaWiFS satellite sensors for six days during the particular sampling period. These two types of data are critical for quantifying both the temporal and horizontal spatial
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3 2
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Figure 4 BTM radiometer data collected with the system shown in Figure 1(D). The top set of panels shows a time series of spectral upwelling radiance (five wavelengths) from the 14-m spectral radiometer system (April–July, 1999). The bottom set of panels shows spectral water-leaving radiance derived from the BTM radiometers and from the SeaWiFS ocean color satellite for six days of the sampling period.
variability of water clarity, PAR, phytoplankton biomass, and primary productivity.
Toward the Future The last two decades have been marked by the emergence of a host of new optical instruments and systems; their applications are growing at a rapid rate. Looking toward the future, it is anticipated that optical instrumentation will be (1) more capable in spectral resolution, (2) smaller and require less
power, (3) suitable for deployment from a variety of autonomous platforms, (4) designed for ease in telemetering of data for real-time applications, and (5) less costly as demand increases enabling higher volumes of data collection.
See also Absorbance Spectroscopy for Chemical Sensors. Bioluminescence. Estimates of Mixing. IR Radiometers. Optical Particle Characterization. Penetrating Shortwave Radiation.
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INHERENT OPTICAL PROPERTIES AND IRRADIANCE
Further Reading Agrawal YC and Pottsmith HC (1994) Laser diffraction particle sizing in STRESS. Continental Shelf Research 14: 1101--1121. Bartz R, Zaneveld JRV, and Pak H (1978) A transmissometer for profiling and moored observations in water. Ocean Optics V, SPIE 160: 102--108. Booth CR (1976) The design and evaluation of a measurement system for photsynthetically active quantum scalar irradiance. Limnology and Oceanography 19: 326--335. Chang GC and Dickey TD (1999) Partitioning in situ total spectral absorption by use of moored spectral absorption and attenuation meters. Applied Optics 38: 3876--3887. Chang GC and Dickey TD (2001) Optical and physical variability on time-scales from minutes to the seasonal cycle on the New England continental shelf: July 1996– June 1997. Journal of Geophysical Research 106: 9435--9453. Chang GC, Dickey TD, and Williams AJ 3rd (2001) Sediment resuspension over a continental shelf during Hurricanes Edouard and Hortense. Journal of Geophysical Research 106: 9517--9531. Dana DR, Maffione RA, and Coenen PE (1998) A new in situ instrument for measuring the backward scattering and absorption coefficients simultaneously. Ocean Optics XIV 1: 1--8. Dickey T (1991) Concurrent high resolution physical and bio-optical measurements in the upper ocean and their applications. Reviews in Geophysics 29: 383--413. Dickey T, Granata T, Marra J, et al. (1993) Seasonal variability of bio-optical and physical properties in the Sargasso Sea. Journal of Geophysical Research 98: 865--898. Dickey T, Marra J, Stramska M, et al. (1994) Bio-optical and physical variability in the sub-arctic North Atlantic Ocean during the spring of 1989. Journal of Geophysical Research 99: 22541--22556. Dickey T, Frye D, Jannasch H, et al. (1998) Initial results from the Bermuda Testbed Mooring Program. Deep-Sea Research I 45: 771--794. Dickey T, Marra J, Weller R, et al. (1998) Time-series of bio-optical and physical properties in the Arabian Sea: October 1994–October 1995. Deep-Sea Research II 45: 2001--2025.
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Dickey TD, Chang GC, Agrawal YC, Williams AJ 3rd, and Hill PS (1998) Sediment resuspension in the wakes of Hurricanes Edouard and Hortense. Geophysical Research Letters 25: 3533--3536. Dickey T, Zedler S, Frye D, et al. (2001) Physical and biogeochemical variability from hours to years at the Bermuda Testbed mooring site: June 1994–March 1998. Deep-Sea Research. Foley D, Dickey T, McPhaden M, et al. (1998) Longwaves and primary productivity variations in the equatorial Pacific at 01, 1401W February 1992–March 1993. Deep-Sea Research II 44: 1801--1826. Gentien P, Lunven M, Lehaitre M, and Duvent JL (1995) In situ depth profiling of particle sizes. Deep-Sea Research 42: 1297--1312. Griffiths G, Knap A, and Dickey T (1999) Autosub experiment near Bermuda. Sea Technology, December. Jerlov NG (1976) Marine Optics. Amsterdam: Elsevier. Kirk JTO (1994) Light and Photosynthesis in Aquatic Ecosystems. Cambridge: Cambridge University Press. Mobley CD (1994) Light and Water: Radiative Transfer in Natural Waters. San Diego: Academic Press. Moore CC, Zaneveld JRV, and Kitchen JC (1992) Preliminary results from in situ spectral absorption meter data. Ocean Optics XI, SPIE 1750: 330--337. O’Reilly JE, Maritorena S, Mitchell BG, et al. (1998) Ocean color chlorophyll algorithms for SeaWiFS. Journal of Geophysical Research 103: 24937--24953. Petrenko AA, Jones BH, Dickey TD, Le Haitre M, and Moore C (1997) Effects of a sewage plume on the biology, optical characteristics and particle size distributions of coastal waters. Journal of Geophysical Research 102: 25 061--25 071. Petrenko AA, Jones BH, and Dickey TD (1998) Shape and near-field dilution of the Sand Island sewage plume: observations compared to model results. Journal of Hydraulic Engineering 124: 565--571. Petzold TJ (1972) Volume scattering functions for selected waters, Scripps Institution of Oceanography Reference 72–78. La Jolla, California: Scripps Institution of Oceanography. Smith RC, Booth CR, and Star JL (1984) Oceanographic bio-optical profiling system. Applied Optics 23: 2791--2797. Spinrad RW, Carder KL, and Perry MJ (1994) Ocean Optics. Oxford: Oxford University Press.
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INTERNAL TIDAL MIXING W. Munk, University of California San Diego, La Jolla, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1323–1327, & 2001, Elsevier Ltd.
Introduction Any meaningful attempt towards understanding how the ocean works has to include an allowance for mixing processes. A great deal is known about the transport of heat and solutes by molecular processes in laboratory-scale experiments. Quite naturally, at the dawn of ocean science, these concepts were borrowed to speculate about the oceans. On that basis it was suggested by Zoeppritz that the temperature structure T(z) at 1000 m depth could be interpreted in terms of the time history T(t) at the surface some 10 million years earlier. It would indeed be nice if past climate could be inferred in this simple manner. Once it was realized that molecular diffusion failed miserably to account for the transports of heat and salt, generations of oceanographers attempted to patch up the situation by replacing the molecular coefficients of conductivity and diffusivity by enormously larger eddy coefficients, but leaving the governing laws (equations) unchanged. By an arbitrary choice of the magnitude of the coefficients it was generally possible to achieve a satisfactory (to the author) agreement between theory and observation, a result aided by the uncertainty of the measurements. But the clear danger signal was there; each experiment, each process required a different set of values. The present generation of oceanographers has come to terms with the need for understanding the mixing processes, just as they had come to terms some years ago with the need for understanding how ocean currents, ocean waves etc. are generated. Wind mixing and tidal mixing are very different processes. And the widespread use of parametrization of mixing processes will not succeed unless there is an underlying understanding of what is being parametrized. Here we have come a long way and have a long way to go.
Stirring and Mixing In a Newtonian fluid the down-gradient flux of a quantity y is given by Fy ¼ k dy=dx
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½1
It is the very smallness of the molecular diffusivity k which requires large gradients dy/dx in order to attain significant fluxes Fy . The fundamental distinction between stirring and mixing was first made in 1948. Stirring produces the gradients whereas molecular mixing reduces the gradients. For the purpose of this article stirring and mixing are both included in the discussion of the contribution of internal tides to mixing processes. For an ocean in steady-state there needs to be an overall balance between the generation and dissipation of mean-square gradients. Nearly all of ocean dynamics deals with processes that generate gradients. This takes place over a wide variety of scales, all the way up to the scales of ocean basins. Dissipation takes place on the ‘microscale,’ i.e., millimeters to centimeters. This is the scale which includes the dominant contributions to the gradient spectrum. A further increase in the spatial resolution of the measurements does not lead to a significant increase in the measured mean-square gradients.
The Battle for Spatial Resolution It is very difficult to attain a quantitative measure of mixing in the turbulent ocean interior. The problem is the need for very high spatial resolution. An eddy coefficient can be estimated as follows: kmolecular rmsðdy=dxÞ ¼ keddy meanðdy=dxÞ
½2
(The subscript ‘molecular’ is introduced here to emphasize the distinction.) The ratio: mean square gradient/square mean gradient (the ‘Cox number’) has been used to estimate the ratio of the eddy coefficient to the molecular coefficient. A typical value away from boundaries is keddy ¼ 105 m2s1, two orders of magnitude in excess of the molecular coefficient. Achieving the required resolution has been a major accomplishment; but there are many problems with the measurements, and even more problems with the interpretation of the measurements along the lines of eqn[2]. It was only with the confirmation by a tracer release experiment that the community has come to accept the value kpelagic ¼ 105 m2s1 for the eddy diffusivity in the interior ocean, away from rough topography. There is of course considerable variability from place to place, but the surprising finding is not how large this variability is but how small it is.
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INTERNAL TIDAL MIXING
Maintaining the Stratification A quite different estimate of eddy diffusivity associated with pelagic mixing can be made from the following considerations. Bottom water is formed in the winter by convective overturning in just a few places: the Greenland Sea, the Labrador Sea and along the Antarctic continent. The formation is estimated at Q ¼ 25 Sverdrups (25 106 m3s1). This would fill the oceans with cold water in a few thousand years. The reason this does not happen is that turbulent diffusion downward from the warm surface balances the upwelling of cold water. With reasonable assumptions this leads to an estimate of eddy diffusivity kstratification ¼ 104 m2 s1, ten times the measured pelagic value. Measurements near topography do indeed give high diffusivities, orders of magnitude above the pelagic value. One simple interpretation is that there are concentrated areas of mixing (just as there are concentrated areas of bottom water formation) from which the water masses (but not the turbulence) are exported into the interior ocean. We can ask the question whether the global stratification can be maintained by vertical mixing in 10% (say) of the ocean volume with an average diffusivity 100 times the pelagic value? The work done against buoyancy by turbulent mixing in a stratified fluid can be written eb ¼ kðg=rð dr=dzÞÞ ¼ kN2 W kg1
½3
where N is the buoyancy frequency. Only a fraction g (called the ‘mixing efficiency’) of the work goes into increasing potential energy (the rest goes into joule heat). A typical value is g ¼ 0.2. The total work per unit area is etotal ¼ eb/g. For the world ocean of area A, the total work done is ð ½4 D ¼ A retotal dz ¼ gg 1kADr W where Dr ¼ 1 kg m3 is taken as the difference between surface and bottom density. Then for A ¼ 3.6 1014 m2 and k ¼ kstratification ¼ 104 m2 s1, one has D ¼ 2 TW (1 terawatt ¼ 1012 W) for the power required to maintain the global stratification in the face of 25 Sverdrups of bottom water formation. To maintain the pelagic turbulence requires only 0.2 TW.
Tidal Dissipation: The Astronomic Evidence It is interesting to compare these numbers with the dissipation of tidal energy. We know this number with remarkable accuracy to be 2.570.1 TW for
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the principal lunar tide (M2); it is obtained from the measured rate of 3.8270.07 cm y1 at which the Moon is moving away from the Earth. For all solar and lunar tides the dissipation is 3.7 TW, but with lesser certainty. We note that the tidal dissipation is of the same magnitude as the 2 TW required for maintaining the ocean stratification. Is this an accident? The astronomic evidence tells us nothing about how and where the dissipation takes place. Allowing for dissipation in the solid Earth and atmosphere leaves 3.4 TW to be dissipated somewhere somehow in the ocean. Ever since it was estimated in 1919 that the dissipation in the Irish Sea is at 0.060 TW, the traditional sink has been in the turbulent bottom boundary layers of marginal seas, about 60 Irish Seas for the world. And before the astronomic estimates settled down to their present value, the ocean estimates kept rising and falling with the astronomic estimates.
Boundary Layer Dissipation versus Scatter When we speak of tides we usually refer to surface (barotropic) tides which have a nearly uniform current velocity from top to bottom, and a maximum vertical displacement at the surface. However, there is also a class of internal (baroclinic) tides with velocities that vary with depth and with maximum displacements in the interior. A surface (barotropic) tide has essentially no shear in the interior ocean, There is shear near the bottom boundary, but the barotropic tidal velocities are so low in the deep ocean that the dissipation is negligible. In shallow seas the barotropic tidal currents are amplified, and the dissipation (proportional to the current cubed) is greatly amplified. This is the basis on which the global tidal dissipation has been attributed to the marginal seas. Internal tides are part of a larger class of internal waves with frequencies other than tidal frequencies. A possible mechanism of tidal dissipation is the scattering of surface tides into internal tides, with subsequent transfer of energy into the broad spectrum of internal waves, and finally into turbulent dissipation: surface tides-internal tides-internal waves-turbulence. What is required at the second stage is some nonlinear frequency splitting which converts the low-frequency tidal line spectrum into a closely packed high-frequency continuum that resembles the observed internal wave spectrum. The final step is associated with the fact that the internal wave spectrum is at or near instability in the
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Richardson sense: the means-square shear is roughly 4 N2 (N is the buoyancy or Brunt-Va¨isa¨la¨ frequency). Scattering of surface tides into internal tides can take place along wavy bottoms. A second possibility is scattering along submarine ridges. An acoustic tomography experiment north at Hawaii detected internal waves of tidal frequency radiating northward.
Satellite Altimetry to the Rescue A subsequent analysis of satellite altimetry clearly showed internal tides emanating from the Hawaii submarine ridge. This is shown in Figure 1. The radiated power was estimated at 0.015 TW. So 14 Hawaiian chains will radiate 0.2 TW, enough to power the pelagic mixing associated with kpelagic ¼ 105 m2 s1. The discovery of internal tide signatures by means of satellite altimetry was an altogether unexpected dividend from a technology that
has revolutionized tidal analysis. The global sampling of surface elevation has introduced a new element into a subject that had gone to bed (in the opinion of some) with the work of Victorian mathematicians. With the Laplace tide equation as a guide for the tidal response of a nondissipative ocean, the assimilation of TOPEX/POSEIDON altimetry can lead to estimates of where one needs to introduce dissipation for agreement with the satellite data. The most recent result allocates roughly 1 TW to the open ocean, mostly over rough terrain. The tentative conclusion is that tidal dissipation is a significant factor in open ocean turbulent mixing. Supporting evidence comes from the measurements of tracer dispersion and microstructure in the Brazil Basin. Diffusivities of k ¼ 2 104 4 104m2 s1 at an elevation of 500 m above the abyssal hills of the Mid-Atlantic Ridge, increasing to 10 104 m2 s1
Figure 1 Surface manifestation of internal M2 tides emanating from the Hawaii’an Island Chain. (Reproduced from Ray and Mitchum 1997.) The wiggly curves show the amplitudes along the ascending orbits of TOPEX/POSEIDON, with positive elevations on the north side. The dashed lines are the inferred crests of the mode 1 component of internal tides. Background shading corresponds to bathymetry, with darker areas denoting shallower water. The triangle to the north east shows the position of the tomographic array. Adapted from Ray & Mitchum (1997).
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INTERNAL TIDAL MIXING
near the bottom have been obtained. Perhaps the most important result is that over a period of a month the diffusivities vary by a factor of two, with the large values occurring at spring tide and the small values at neap tide.
Discussion There is more than enough tidal dissipation to feed the measured pelagic turbulence associated with kpelagic ¼ 105 m2 s1. With regard to the larger value kstratification ¼ 104 m2 s1, the present best estimates would suggest that tidal dissipation could power half the turbulence needed to account for the observed ocean stratification. Figure 2 attempts an allocation of tidal energy flux, but there are many uncertainties, some by factors of two or more. The assumed onedimensional balance between upward advection and
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downward diffusion as a measure of kstratification is itself somewhat uncertain. However, the present conclusion is that tidal dissipation is a significant, possibly dominant factor driving mixing in the ocean interior. The combination of detailed in situ measurements of turbulent mixing subject to a global lid on available tidal energy has led to giant strides towards a meaningful parametrization of ocean mixing, whether or not tidally induced. The conversion of wind energy to turbulent mixing plays a major role, particularly in the upper oceans. In Figure 2, equal weight has arbitrarily been assigned to tides and winds. We shall have to await the outcome of this competition.
See also Dispersion and Diffusion in the Deep Ocean. Internal Tides. Internal Waves. Tides. Turbulence in the Benthic Boundary Layer.
Further Reading
2
3
Figure 2 Sketch of proposed flux of tidal energy (modified from Munk and Wunsch, 1997). The traditional sink is in the turbulent boundary layer of marginal seas. Scattering into internal tides over ocean ridges (by the equivalent of 14 Hawaii’s) and subsequent degradation into the internal wave continuum feeds the pelagic turbulence at a level consistent with kpelagic ¼ 105 m2 s1. Most of the ocean mixing is associated with a few concentrated areas of surface to internal mode convergence over regions of extreme bottom roughness and with severe wind events. Light lines represent speculation with no observational support.
Cartwright DE (1999) Tides; a Scientific History. Cambridge: Cambridge University Press. Dushaw BD, Cornuelle BD, Worcester PF, Howe BM, and Luther DS (1995) Barotropic and baroclinic tides in the central North Pacific Ocean determined from longrange reciprocal acoustic transmission. Journal of Physical Oceanography 25: 631--647. Eckart C (1948) An analysis of the stirring and mixing processes in incompressible fluids. Journal of Marine Research 7: 265--275. Gregg MC (1988) Diapycnal mixing in the thermocline; a review. Journal of Geophysical Research 92: 5249--5286. Ledwell JR, Watson AJ, and Law CS (1993) Evidence for slow mixing across the pycnocline from an open-ocean tracer release experiment. Nature 364: 701--703. Munk W (1997) Once again: once again-tidal friction. Progress in Oceanography 40: 7--36. Munk W and Wunsch C (1997) The Moon, of course. Oceanography 10: 132--134. Munk W and Wunsch C (1998) Abyssal recipes II: energetics of tidal and wind mixing. Deep-Sea Research I 45: 1977--2010. Ray RD and Mitchum GT (1997) Surface manifestation of internal tides in the deep ocean: observations from altimetry and island gauges. Progress in Oceanography 40: 135--162. Taylor GI (1919) Tidal friction in the Irish Sea. Philosophical Transactions of the Royal Society A 220: 1--93.
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INTERNAL TIDES R. D. Ray, NASA Goddard Space Flight Center, Greenbelt, MD, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1327–1335, & 2001, Elsevier Ltd.
Introduction Oceanic internal tides are internal waves with tidal periodicities. They are ubiquitous throughout the ocean, although generally more pronounced near large bathymetric features such as mid-ocean ridges and continental slopes. The internal vertical displacements associated with these waves can be extraordinarily large. Near some shelf breaks where the surface tides are strong, internal displacements (e.g., of an isothermal surface) can exceed 200 m. Displacements of 10 m in the open ocean are not uncommon. The associated current velocities are usually comparable to or larger than the currents of the surface tide. Internal tides can occasionally generate packets of internal solitons which are detectable in remote sensing imagery. Other common nonlinear features are generation of higher harmonics (e.g., 6 h waves) and wave breaking. Internal tides are known to be an important energy source for mixing of shelf waters. Recent research suggests that they may also be a significant energy source for deep-ocean mixing. Internal tides were first recognized in the early part of the twentieth century, yet as late as the 1950s arguments were still being waged over what causes them. Their wavelengths, generally shorter than 200 km, are poorly matched to the planetary-scale astronomical tidal potential, so the generation mechanism for surface tides appears inapplicable. Various theories invoking hypothetical resonances at inertial latitudes (where tidal and Coriolis frequencies are equal) were put forward, but they are not compelling, not least because the inertial latitude for the dominant tide M2 is in the far polar latitudes (74.51). The now accepted explanation for internal tides is that they are generated by the interaction of the barotropic surface tide with bottom topography. As the tide sweeps stratified water over topographic features, it disrupts normal (equilibrium) isopycnal layers, setting up pressure gradients that induce secondary internal motions at the same frequency as the tide. Since internal tides are a special kind of internal wave, much of our knowledge of internal waves is immediately applicable. For example, an internal
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tide always displays current shear – i.e., the associated horizontal current velocities change with depth – whereas the surface tide’s horizontal current is independent of depth. And like other internal waves in smoothly varying density stratification, an internal tide displays the seemingly odd property that its group velocity is in the same vertical plane but perpendicular to its phase velocity. The fundamental properties of internal tides, including whether or not they even exist, are controlled by the relative magnitudes of three basic frequencies: the tidal frequency o, the local Brunt-Va¨isa¨la¨ or buoyancy frequency N, and the local Coriolis frequency f. Depending on which of these frequencies is highest and which lowest, internal tides may propagate freely away from their generation point, they may be reflected in some manner, or they may be evanescent. For midlatitude semidiurnal tides, typically f oooN, a regime allowing free propagation. Given that the generation and propagation of internal tides depend strongly on the stratification, it is not surprising that most observations have found internal tides to be highly variable, sometimes with pronounced seasonal variations. In some places they appear only during spring tides (when solar and lunar tides are at maximum). In other places they appear randomly intermittent, evident for several days and then disappearing. Some observations, primarily from the open ocean, have revealed a component that does remain temporally coherent with the astronomical potential (see below), but the dominant characteristic of internal tides in most regions is one of incoherence, both spatially and temporally.
Modes and Beams Two complementary dynamical frameworks are used for analyzing internal tides: decomposition into vertical modes and propagation along characteristics. Generally, the latter description is more useful near generation points, and the modal description more useful elsewhere, but in any particular situation one or the other approach may be advantageous. Both approaches require knowledge of the stratification, usually parameterized by the buoyancy frequency N. This is the frequency with which a vertically displaced fluid element would oscillate because of restoring buoyancy forces. It is given by N ¼ Oðgr1 @r=@zÞ, where g is the acceleration of gravity and r is the average potential density, a
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INTERNAL TIDES
function of position and depth. The Coriolis frequency f ¼ 2Osin f, where f is the latitude and O is the Earth’s sidereal rotation frequency (7.2921 105s1).
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Vertical displacement Sea surface
Modes
Mode 1
The governing dynamical equations for internal tides, under linear, hydrostatic, inviscid, Boussinesq and flat-bottom assumptions, and neglecting the horizonal currents, may be solved by separation of variables. The equation for the vertical displacement leads to an eigenvector problem with eigenvalue a2n :
Mode 3
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q Gn ðzÞ þ a2n N 2 ðzÞGn ðzÞ ¼ 0 qz2
Mode 5
The eigenvectors Gn ðzÞ, ordered so that anþ1 > an , provide a complete, orthogonal basis for the internal vertical displacements. The corresponding equations for horizontal dependence yield expressions for the horizontal wave number k, phase velocity cp ¼ o=k, and group velocity cg ¼ do=dk in terms of an : k2 ¼ a2n o2 f 2
c2p ¼ o2 = a2n o2 f 2 cg ¼ cp a2n
Seafloor Figure 1 Baroclinic displacement modes 1, 3, and 5, computed for a buoyancy frequency profile from the deep ocean.
1
If N is taken constant, then Gn can be found analytically: Gn ðzÞ ¼ asinðnpz=DÞ for an ocean depth D. If N is taken more representative of deep-ocean conditions, with a peak at the pycnocline, then the oscillations in Gn are shifted upward (Figure 1). Notice that the displacements are small or zero at the top and bottom of the water column, and that each Gn has n 1 crossings of the origin. Horizontal velocity modes are given by dGn =dz, and they have n crossings and hence nonzero shear for all modes. Most observations of internal tides (except those very near the generation point) are adequately described by a superposition of a few low order modes. In the deep ocean typical phase speeds cp are of order 3 m s1 for n ¼ 1. Corresponding wavelengths l ¼ 2p=k are between 100 and 200 km. Higher order modes have speeds and wavelengths given roughly by cp =n and l=n, respectively. On continental shelves both speeds and wavelengths may be an order of magnitude smaller. These values are for semidiurnal tides; wavelengths of diurnal tides are approximately twice as large. From the above expressions for k and cp it is apparent that internal tides cannot freely propagate unless o > f . They are ‘evanescent’ (exponentially
damped)’ polewards of the critical latitudes where o ¼ f . Freely propagating waves for diurnal tides are therefore confined to the region between latitudes 7301. (In fact, unambiguous observations of diurnal tides are fairly rare, but this is partly due to relatively weak barotropic forcing and higher background noise levels.) Beams
A complementary approach to modal analyses stems from the equation for the two-dimensional stream function, which is hyperbolic in spatial coordinates and may therefore be solved by the method of characteristics. The resulting solution consists of narrow beams of intense motion embedded in an otherwise resting ocean. The group velocity, and hence the energy propagation, follow the characteristics, which are along lines of slope: 2 1=2 o f2 c ¼ tan y ¼ 7 N 2 o2 From a given internal tide generation point, energy thus propagates along beams at the angle y relative to
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horizontal, the angle depending (for a given f and N) only on the tidal frequency. Well defined beams comprise a large number of modes, with modal cancellations occurring outside the allowed beam. A numerical example of beam-like propagation from a shelf break is shown in Figure 2. Generation of internal tides is apparently especially efficient when the seafloor slopes at precisely the critical value c. Barotropic flow is then coincident with the motion plane for free internal waves, resulting in near-resonant conditions in which even quite small surface tides can generate internal tides. With nominal values of NB50 cpd, oB2 cpd, f B0:6 cpd, then y is 21. Continental slopes commonly exceed this, so c would be attained near the shelf break, as depicted in Figure 2. When an internal wave is reflected from the ocean bottom or ocean surface, energy propagation is still confined to the angle y, which makes for a curious variation on the usual laws of reflection. If the wave 0 4 80 4 80 4 80 4 80 4
80 4
is incident upon bathymetry that is steeper than y (supercritical case), energy is reflected backwards into deeper water. If the bathymetry is less steep (subcritical case), energy is reflected forward toward shallower water (see Figure 3). Ocean observations of this behavior are not easy to obtain, since mooring instruments must be precisely placed (depending on the ambient N); yet measurements of internal tides in the Bay of Biscay have not only observed the downward energy propagation from the generation point, but also the subsequent reflection from the ocean bottom. In the Bay of Biscay, as in most places, N diminishes with depth, so y grows larger and the beams become steeper as they approach the bottom (Figure 2 and Figure 3 are drawn for constant N.) With such reflection properties, internal waves incident on a subcritical sloping bottom will be focused into the shallows (as in Figure 3A), with energy density correspondingly intensified. The same mechanism tends to trap internal wave energy within steep 80 4 8 0 4 8 0 4 8 0 4 8 0 4 8 0 4
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Figure 2 Theoretical internal wave beam from a shelf edge in a constant N ocean. (A) vertical displacements following beam at constant slope tan y; (B) phase contours (in degrees) of the vertical displacements relative to the surface tide. Notice how phase propagation is at right angles to the beam; i.e., the phase velocity is perpendicular to group velocity (and to the direction of energy propagation). (Reproduced with permission from Prinsenberg SJ and Rattray M (1975) Effects of continental slope and variable BruntVa¨isa¨la¨ frequency on the coastal generation of internal tides. Deep Sea Research 22: 251–263.)
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downward edge the tide has generated packets of internal solitons. These solitary waves, starting from an initial 7 m depth, have extraordinarily nonlinear isotherm displacements of 25 m. Although usually less dramatic, this phenomenon is not uncommon in high tide regions on a continental shelf. The internal tide, generated at the shelf break, propagates shorewards and becomes progressively more nonlinear until the bore disintegrates into a group of solitons, the leading one usually of largest amplitide.
(A)
Moored Current Meters
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Figure 3 Successive reflections of an internal wave along (A) a subcritical seafloor and (B) a supercritical seafloor. Arrows denote direction of energy propagation.
(supercritical) canyons, where the canyon sides reflect energy ever deeper, focusing it toward the canyon floor. If the floor is subcritical, then energy is further focused toward the canyon head. Intense internal tide currents and large kinetic energy densities have indeed been observed in canyons, and especially near canyon heads. In the presence of internal viscosity or other dissipative mechanisms, internal tidal beams widen. Because group velocities are smaller and decay scales shorter for higher order modes, beams tend to disintegrate rapidly into the few low order modes that are most commonly observed.
Observations Internal tides have been observed with a great multitude of instruments and technique, both in situ and remote. Four distinctly different examples are given here which serve to highlight a number of characteristic features of internal tides. Except for the first example, emphasis is given to deep-sea tides. Vertical Profilers
Vertical profilers, ranging from echo sounders to repeated hydrographic casts to special yo-yo instruments, provide some of the clearest pictures of internal tides. An especially dramatic example from the continental shelf off Oregon is shown in Figure 4. It shows a clear semidiurnal signal in the isotherm displacements, somewhat distorted into a bore-like shape (akin to the nonlinear distortion seen in shoaling surface tides in very shallow water). Along its
Because of their widespread deployments, current meters provide perhaps the most common for observing, or at least detecting, internal tides, especially in the open ocean. Sufficient vertical sampling is required for decoupling the internal modes from the surface tide (and unfortunately sufficient sampling is not common). Figure 5 is an example of marginally adequate vertical sampling; it shows tidal current estimates extracted from moored meters near 1101W on the Pacific equator. Estimates are given for each of 10 months, at 10 depths throughout the water column. The current ellipses are fairly uniform below 1000 m; these depths are dominated by the stable, depth-independent currents of the surface tide. In contrast, large temporal variation, and occasionally much larger amplitudes, are evident in the shallower estimates; in these depths, where the buoyancy frequency (and its change) is maximum, the tidal signal is dominated by the internal tide. Modal analysis reveals that the internal tide is essentially random, isotropic, and without a dominant mode for these 10 months. Such observations are characteristic of in situ observations of internal tides; but in a few locations in the deep ocean, a component of the internal tide has been observed that is not so variable and that maintains phase lock with the astronomical tide. The famous MODE experiment in the western Atlantic found that approximately 50% of the internal tide variance was temporally coherent with the astronomical tide. Such observations imply a nearly constant ocean stratification, at least to the extent that it determines generation and propagation properties. Satellite Altimetry
Recently satellite altimetry has been shown capable of providing a near-global view of the coherent component of internal tides. It does this by detecting the very small surface displacements associated with internal tides. These are given roughly by the tide’s internal displacements scaled by Dr=r, the fractional difference in water density, typically of order 0.2%,
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Temperature (˚C) 18
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Figure 4 Color contour time series of temperature profiles from the surface to 35 m depth, obtained by repeated 80 s raising and free-falling of the loose-tethered microstructure profiler, deployed offshore Tillamook, Oregon in October 1995. Top: the semidiurnal internal tide displacement (most clearly seen along the yellow 13.81C isotherm) for a 24 h period. Bottom: a zoom view of a 1.7 h period showing the start of the first soliton displacements. The solitons are separated by roughly 10 min. (Reproduced with permission from Stanton TP and Ostrovsky LA (1998) Observations of highly nonlinear internal solitons over the continental shelf. Geophysical Research Letters 25: 1695–1698.)
thus implying surface displacements of a few centimeters for internal displacement of tens of meters. Altimetry detects such small waves as modulations (with wavelengths 100–200 km for internal mode 1)
of the surface tide as estimated along satellite tracks. Because tides can be estimated from altimeter data only by gathering multi-year time series of elevations at a particular site, only the coherent component of
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INTERNAL TIDES
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2 cm/s +
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Figure 5 M2 tidal current ellipses by month at each of 10 depths obtained from moored current meters near 01, 1101W. Scale bar for velocity is at upper right. Each ellipse indicates how the direction and magnitude of the horizontal current velocity evolves over one tidal cycle. All ellipses are polarized clockwise except those marked with a plus sign. (Reproduced with permission from Weisberg RH, Halpern D, Tang T, Hwang SM (1987) M2 tidal currents in the eastern equatorial Pacific Ocean. Journal of Geophysical Research 92: 3821–3826.)
the internal tide which maintains phase lock with the surface tide is capable of being detected. Figure 6 gives an example of the first detection of such waves, near the Hawaiian Ridge. The waves are roughly 5 cm amplitude near the ridge and decay slowly with distance, but are still detectable 1000 km away. Phase estimates (not shown) reveal clearly that the waves are propagating away from the ridge. Evidently they are created by the barotropic tide striking the ridge (at nearly right angle from the north) and generating an internal tide that propagates both northwards and southwards. The picture reveals three important aspects of deep-ocean internal tides: (1) that in some locations they maintain temporal coherence over several years, thus allowing altimetry to measure them, (2) that they maintain spatial coherence over a wide area, and (3) that they are capable of propagating hundreds to thousands of kilometers before being dissipated. All three aspects contrast sharply with the usual picture of incoherence obtained from in situ observations.
Waves similar to those in Figure 6 have been detected in many regions throughout the global ocean. However, altimetry is incapable of detecting internal tides in a region where they are temporally incoherent. Such is apparently the case, for example, off the northwest European shelf, a region known for some of the largest internal tides in the world, but where the coherent signals in altimeter data are extremely weak. Internal tide studies with satellite altimetry are relatively new, and further work should reveal new facets from a global perspective. Acoustic Methods
A second example of a powerful, but unconventional, technique for studying coherent internal tides is acoustic tomography. Differences in two-way acoustic travel times between reciprocal transceivers are sensitive to barotropic tidal currents within the acoustic path. Similarly, since vertical isotherm displacements perturb the sound speed
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10 cm
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Figure 6 Mean elevations at the sea surface of internal tides near Hawaii, deduced from altimeter measurements of the Topex/ Poseidon satellite. Positive values (north of the trackline) indicate that the internal tide’s surface elevation is in phase with the barotropic tide’s elevation. Scale bar for elevations at upper left. Background shading corresponds to bathymetry, with darker shading denoting shallower depths and the main axis of the Hawaiian Ridge. Only internal tides that are coherent with the surface tide over the entire measurement period (here 3.5 years) can be detected in this manner.
within the path, the sums (or averages) of the travel times are sensitive to the internal tide. From a sufficiently long time series the mean tidal characteristics along a given path can be determined. The seemingly coarse spatial resolution is actually an advantage, because it suppresses short-scale internal waves and other noise that typically plague current meter measurements. And, in fact, an array of acoustic transceivers can act as a very sensitive directional antenna for spatially coherent internal tides. Such an array in the central Pacific, consisting of acoustic paths roughly 1000 km long and located just north of the area shown in Figure 6, has measured the same coherent internal tide field seen in the altimetry and indicates that the primary source is the Hawaiian Ridge, even at that great distance.
Implications for Energetics and Mixing Internal tides are an important energy source for vertical mixing, especially in coastal waters where they help maintain nutrient fluxes from deep water to euphotic zones on the shelf. A good example is the
Scotian shelf off Nova Scotia where internal tides are responsible for a strip of enhanced concentrations of nutrients and biomass along the shelf break. During each tidal cycle one or two strong (50 m) internal solitons (compare Figure 4) are generated near the shelf edge, moving shoreward but dissipating rapidly, possibly within 10 km. Estimated energy fluxes of 500 W m1 appear more than adequate to maintain observed nutrient supply to the mixed layer. Similar mixing mechanisms have been observed in the Celtic Sea and elsewhere. In the open ocean it seems reasonable that internal tides dissipate by transferring energy into the internal wave continuum or by directly generating pelagic turbulence, but the associated energy fluxes, and even the dominant mechanisms, are unclear. Nonlinearity is a common feature of internal tides (e.g., occurrences of higher harmonics), so ‘diffusion’ into the continuum is conceivable via nonlinear (resonant triad) interactions, but the evidence for this is so far more anecdotal than convincing. Bottom scattering of low mode tides into higher modes may play a role, as well as wave reflections off sloping bottoms, which tend to intensify kinetic energy densities and may lead to shear instabilities and wave breaking.
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INTERNAL TIDES
The traditional view is that both the internal wave continuum and the pelagic turbulence and mixing are maintained by nontidal mechanisms such as wind generation; whether internal tides play a major or minor role in this is not resolved. At a minimum, improved quantitative estimates are needed for the global internal tide energy budget. The internal tide energy budget also has a bearing on a longstanding geophysical problem: finding the energy sink for the global surface tide. If the generation/dissipation rate for internal tides is sufficiently large, then internal tide generation conceivably supplements the traditional sink of botton friction in shallow seas. Dissipation rates for the surface tide are well determined by space geodesy (e.g., lunar laser ranging) at 3.7 TW, with 2.5 TW for the principal tide M2. How much of this is accounted for by conversion into internal tides is not well determined; published estimates range from o100 GW (0.1 TW) to >1 TW. There is fairly wide agreement that generation of internal tides at continental slopes provides a fairly small energy sink. Both models and measurements suggest that typical energy fluxes at shelf breaks are of order 100 W m1, leading to a global total of order 15 GW. This is perhaps an underestimate, because it may not fully account for shelf canyons and other three-dimensional features, but the order of magnitude seems reliable. Internal tide generation by deep-ocean topography, however, may be far more important. Recent research based on global tide models as well as on empirical estimates of tidal dissipation deduced from satellite altimetry suggests that generation of internal tides by deep-sea ridges and seamounts could account for 1 TW of tidal power. Refining such estimates, and understanding the role that internal tides play in
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generation of the background internal wave continuum, in vertical mixing, and in maintenance of the abyssal stratification, are some of the outstanding issues of current research.
See also Acoustics in Marine Sediments. Internal Tidal Mixing. Internal Waves. Tides.
Further Reading Baines PG (1986) Internal tides, internal waves, and nearinertial motions. In: Mooers C (ed.) Baroclinic Processes on Continental Shelves. Washington: American Geophysical Union. Dushaw BD, Cornuelle BD, Worcester PF, Howe BM, and Luther DS (1995) Barotropic and baroclinic tides in the central North Pacific Ocean determined from longrange reciprocal acoustic transmissions. Journal of Physical Oceanography 25: 631--647. Hendershott MC (1981) Long waves and ocean tides. In: Warren BA and Wunsch C (eds.) Evolution of Physical Oceanography. Cambridge: MIT Press. Huthnance JM (1989) Internal tides and waves near the continental shelf edge. Geophysical and Astrophysical Fluid Dynamics 48: 81--106. Mun WH (1997) Once again: once again – tidal friction. Progress in Oceanography 40: 7--35. Ray RD and Mitchum GT (1997) Surface manifestation of internal tides in the deep ocean: observations from altimetry and island gauges. Progress in Oceanography 40: 135--162. Wunsch C (1975) Internal tides in the ocean. Reviews of Geophysics and Space Physics 13: 167--182.
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INTERNAL WAVES C. Garrett, University of Victoria, VIC, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1335–1342, & 2001, Elsevier Ltd.
Introduction Waves at the sea surface are a matter of common experience. Surface tension is the dominant restoring force for waves with a wavelength less than 17 mm or so; longer waves are more affected by gravity. They have periods up to about 20 s and amplitudes that may be many meters. Given the stable density stratification of the ocean, it is not surprising that there are also ‘internal gravity waves,’ with a water parcel displaced vertically feeding a gravitational restoring force. The wave periods depend on the degree of stratification but may be as short as several minutes and can be long enough that the Coriolis force plays a major role in the dynamics. Vertical displacements are typically of the order of ten meters or so, with horizontal excursions of several hundred meters. The associated horizontal currents are typically several tens of millimeters per second. An interesting difference from the surface wave field is that internal waves always seem to be present, without the intense storms or periods of calm that exist at the surface. The existence of internal waves complicates the mapping of average currents and depths of particular density surfaces. They have also been the objective of intensive military-funded research because of the possibility that wakes of internal waves generated by submarines might be detectable by remote sensing, thus betraying the submarine’s location. More conventional acoustic means of submarine detection are complicated by the deflection of acoustic rays by the rather random variations in sound speed induced by internal waves. In civilian activities, the currents and buoyancy changes associated with internal waves are a matter of concern in offshore oil drilling. Most importantly, perhaps, the current shear of internal waves, including those of tidal frequency, can lead to instability and turbulence, and so the waves are the main agent for vertical mixing in the ocean interior. This mixing plays a major role in determining the strength of ocean circulation, and hence the poleward heat flux and climate. The mixing, along with the associated circulation, also
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provides nutrient fluxes into the sunlit upper ocean where primary biological production occurs. Understanding internal waves is thus of vital importance, particularly since they occur at too small a scale to be treated explicitly in computer models of the ocean. Their effects must be ‘parametrized,’ or represented by formulas that involve only the quantities that are carried in the model. In this respect, internal waves in the ocean are somewhat akin to clouds in the atmosphere – they play a vital, perhaps even controlling, role in global-scale problems. (The atmosphere also has internal waves, of course, which are known to play a major role in redistributing momentum.) This short article will first describe the waves that can occur at sharp density interfaces in clear analogy to waves at the sea surface. This will be followed by a description of the waves that can propagate through a continuously stratified ocean, and a discussion of their generation, evolution, and relationship to ocean mixing.
Interfacial Waves If the ocean consists of an upper layer of density r Dr and thickness h1 above a layer of density r and thickness h2, then waves that have a wavelength much greater than both h1 and h2 travel at a speed [g0 h1h2/(h1 þ h2)]1/2 independent of wavelength, where g0 ¼ gDr/r is known as the ‘reduced gravity’. If h2bh1, this becomes (g0 h1)1/2, in clear analogy to the speed (gh)1/2 for surface waves that are long compared with the water depth h. This formula for the speed of interfacial waves also holds for h2bh1 even if the wavelength is not long compared with h2. The theory behind this requires that the amplitude of the waves is much less than the thickness of the layers. Many observed interfacial waves (Figure 1) violate this assumption and also the requirement that their wavelength is long compared with the layer thicknesses. Finite amplitude is associated with a tendency for waves to steepen, much as in the development of a tidal bore at the sea surface. On the other hand, a horizontal scale that is not very long, compared with at least the thickness of the thinner layer, leads to dispersion, the break-up of a disturbance into waves of different wavelengths traveling at different speeds. Interestingly, these effects can cancel, leading to the possibility of ‘internal solitary waves’, waves of finite amplitude that can be spatially localized and travel without change of shape. They can occur singly, or in groups as in Figure 1.
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Figure 1 A group of interfacial solitary waves generated by tidal flow over the sill at Knight Inlet, British Columbia. (A) The banded surface manifestation. (B) An echo-sounder image of the same waves. Current vectors are also shown. The location and ship track are shown in the bottom inset. The upper left inset of the density (st) profile shows a strongly stratified thin upper layer above a much thicker more homogeneous layer, rather than an ideal two-layer situation. (From Farmer D and Armi L (1999) The generation and trapping of solitary waves over topography. Science 283: 188–190; courtesy of D. Farmer.)
If in a group, the crests pointing away from the thinner layer are sharper than the troughs. Even if they occur at a density interface many meters, or tens of meters, below the surface, they are often visible if the upper layer is turbid, so that the crests appear from above as more opaque tubes than the surrounding water. More frequently they are seen because the associated currents cause visible variations in surface roughness (Figure 1A). (Whether the water is rougher above the crests or troughs of the interfacial waves depends on the relative directions of propagation of the surface waves and the interfacial waves, as can readily be established by considering the interaction in a frame of reference moving with the interfacial waves.) The generation of these packets of internal solitary waves, or trains of waves with similar properties, often occurs when tidal flow over a sill, or off the edge of the continental shelf, leads to a leeward
depression in the interface (as in Figure 1). As the tidal current reverses, this depression propagates back over the sill, or onto the continental shelf, and breaks up into large amplitude interfacial waves. (In the situation shown in Figure 1, interfacial waves have actually formed before the current reversal.) The internal solitary waves, or solitons, typically have periods of tens of minutes. This would appear to be too short for the Earth’s rotation to be a factor, but it does seem that the break-up of an internal tide into internal solitons may be inhibited by rotational effects. While it is generally only the shape of these interfacial waves that propagates, with little net water movement, they can be sufficiently large that they do, in fact, carry water along with them. Remarkable behavior also occurs as the waves approach shore: although they have sharp downward crests offshore where the lower layer is thicker,
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they must switch to having sharp upward crests as the lower layer becomes thinner than the upper layer! They do this only with considerable loss of energy into smaller-amplitude dispersive waves and turbulence. The waves described above have a vertical motion that is maximum at the interface and tends to zero at top and bottom boundaries. The horizontal currents in the two layers are in opposite directions. A sharp interface is not necessary; large-amplitude internal motions can persist even if the density jump is smeared out in the vertical. In the case of smallamplitude waves, the motion can then be thought of as the first vertical mode, made up of propagating waves that reflect off the sea surface and seafloor. We therefore turn next to a discussion of these building blocks.
Internal Waves The Basic Physics
A particle displaced vertically in a continuously stratified fluid experiences a restoring buoyancy force. If a whole vertical fluid column is displaced, the vertical uniformity of the motion means that there is no change in the hydrostatic vertical pressure gradient and the restoring force on each particle is just gravity g times the density perturbation, which is minus the vertical displacement times the vertical density gradient dr/dz. This leads to a simple harmonic oscillator equation for the motion of the column, with the frequency of oscillation given by the ‘buoyancy frequency’ N, where N2 ¼ (g/r)(dr/dz). This frequency is independent of the horizontal scale of the fluid columns, suggesting that the frequency of motions that are wavelike in the horizontal is N, independently of scale. If the fluid columns are now allowed to oscillate obliquely, at an angle y to the vertical, the vertical restoring force is reduced by a factor cos y, as is the component of this force parallel to the motion. The extra factor cos2 y in the oscillator equation then means that the frequency is reduced to N cos y, again independently of the lateral scale. Regarding these motions as waves, it is clear that the motion is transverse, as required for an incompressible fluid with = u ¼ 0. In a rotating world the motion is acted upon by the Coriolis force, so that fluid oscillations in inclined sheets now develop a transverse motion, within the sheet but orthogonal to the motion with no rotation. The relationship between the frequency o of the oscillations now involves the Earth’s rotational frequency. Provided that N is sufficiently greater than the
Coriolis frequency f, which is twice the vertical component of rotation, the connection between frequency o and orientation y becomes eqn [1]. o2 ¼ N 2 cos2 y þ f 2 sin2 y
½1
Equation [2] is an alternative expression, in terms of the wavenumber k ¼ (k, l, m). N 2 k2 þ l2 þ f 2 m2 o ¼ k2 þ l2 þ m2 2
½2
While the frequency can be as high as N when the particle motion is vertical, it cannot be lower than the Coriolis frequency f. In this limit, the particle motion is horizontal in ‘inertial’ circles, expressing the tendency for steady rectilinear motion with respect to a nonrotating reference frame. Any frequency of motion between these limiting frequencies is possible, depending on y, or, equivalently, the ratio of vertical to horizontal wavenumber. Moreover, at any frequency, any wavelength is possible. The group velocity (the velocity with which a wave packet, or energy, propagates) is given by (qo/qk, qo/ ql, qo/qm). For internal waves this is easily shown from (2) to be at right angles to the wavenumber vector k. In other words, energy propagates parallel to the wave crests, rather than at right angles as for surface waves! This remarkable feature can be demonstrated in a laboratory experiment (Figure 2). In terms of vertical propagation, waves with downwards phase propagation have upwards energy flux, and vice versa. Observations
Measurements of the frequency spectra of internal waves can be obtained from analysis of time-series of measurements, at a fixed point, by current meters or by temperature sensors that show changes associated with vertical motion of the temperature-stratified water. Such measurements do show a block of energy at frequencies between f and N, falling off above and below these frequencies and with an energy distribution in between that seems close to o2. For currents there is typically an extra peak (an ‘inertial cusp’) near f, but this is suppressed in temperature data as the near-inertial motions are largely horizontal. Measurements at a single fixed point do not, however, provide information on the wavenumber content, or spatial scales, of the energy at any frequency. For this one needs information from many locations (such as from many current meters on a mooring) or the continuous vertical profile obtainable from an acoustic Doppler current profiler
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INTERNAL WAVES
m
269
=m *
1
cg
cg
10 _2
10
k
k
cg
k
k
cg
Figure 2 Internal waves of a fixed frequency less than N are generated by vertical oscillations of a wavemaker in a tank of stratified fluid. In the photograph the light and dark radial lines are contours of constant density perturbation and so are lines of constant wave phase. The schematic diagram shows the directions of wave phase propagation (k) and group velocity (cg).
(ADCP). Current meter arrays are, of course, limited by the cost and logistics of deploying large numbers of instruments, and moored ADCPs with good resolution have a range much less than the depth of the ocean. Invaluable high-resolution vertical profiles of horizontal currents over the whole ocean depth have been obtained from dropped or lowered profiling current meters that measure the tiny electric potentials generated by movement of conducting sea water in the Earth’s magnetic field, and also by lowered ADCPs. These techniques do not, however, provide much information on the frequency content of the motions. Further information has also come from horizontal tows of sensors, or arrays of sensors, hence mapping horizontal scales though, again, not providing frequency information. Various syntheses of the information from these types of experiments have shown a tendency for energy to be distributed in vertical wavenumber and
f
10
_1
10
N
m
2
1
Figure 3 The so-called ‘Garrett–Munk’ spectrum, giving the distribution of energy in a space defined by frequency o and vertical wavenumber m. The spectrum has a peak at the inertial frequency f and falls off like o2 at higher frequencies up to N. In vertical wavenumber the spectrum is fairly flat at small wavenumbers (large scale), then falls off rapidly to high wavenumber. The scales are as multiples of f for o and in terms of an equivalent vertical mode number (number of halfwavelengths in the ocean depth) for m.
frequency somewhat as shown in Figure 3: the frequency dependence is roughly like o2, and at each frequency there is a tendency for there to be more energy at small vertical wavenumbers (large scales), with a roll off to high wavenumbers with a power law like m2 or m5/2, though the exponent here is certainly not well-established or universal. While a given frequency and vertical wavenumber magnitude are associated with a given magnitude of horizontal wavenumber via [2], the direction of the wavenumber is not specified, either up versus down or in the 3601 available horizontally. It is generally assumed that the energy is horizontally ‘isotropic,’ or distributed evenly among all possible directions of propagation. It is also assumed that there is as much energy propagating up as down, though there is evidence for preferential downward transmission for waves with frequency within about 10% of f at midlatitude. The inertial peak, in fact, deserves special consideration, given its dominance. One interesting aspect is that the current vectors spiral with depth, with a connection between the direction of rotation of the spiral and the direction of energy propagation: in the Northern Hemisphere currents rotate clockwise with time, so that a current vector profile that
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shows increasing clockwise rotation with increasing depth below the sea surface must have phase propagation upward and hence group, and energy, propagation downward. The model spectrum shown in Figure 3 is only a very rough approximation. Observed spectra typically have considerable additional energy at tidal frequencies, with this energy also distributed over various vertical scales.
or standing internal waves behind topographic features (as occurs in the atmosphere). These may propagate upward into the ocean. The relative importance of these different mechanisms on a global basis has not been established, but the current view is that wind and tides are the dominant sources and are of comparable importance. Evolution
Generation
It seems that the inertial peak may be generated by fast-moving storms that set up currents in the upper ocean with the corresponding Coriolis forces unmatched by pressure gradients. Near-inertial motions result, with the current vectors rotating with a frequency close to the local f. The large horizontal scale of these motions means, however, that they experience different f at different latitudes. The current vectors at different latitudes then rotate at different rates, increasing the latitudinal gradients in current vectors and decreasing the horizontal scale. The resulting convergence and vertical motion means that the waves can no longer be purely inertial; they retain their frequency but propagate equatorward to a region where f is smaller. At the same time they develop an increasing vertical group velocity and propagate downward into the ocean. The evolution of this important near-inertial part of the internal wave spectrum is also affected by wave interactions with lower-frequency eddies. Higher-frequency waves may be generated by storms that move more slowly, by turbulence in the surface mixed layer, by subtle interactions between surface waves, or as part of the decay process of ocean eddies. They may also arise from interactions between preexisting internal waves, as will be discussed shortly. The tides are another important source of energy for internal waves observed throughout the ocean, as already discussed for interfacial waves near the sea surface. The barotropic, depth-independent, tidal currents associated with tidal changes in sea level move density-stratified water over topographic features on the seafloor, setting up internal oscillations much as if the topographic features were oscillating wavemakers in an otherwise still ocean. These ‘internal tides’ are mainly at the tidal frequencies, though there may also be energy at multiples of these. Lower-frequency currents in the deep ocean are generally much weaker than tidal currents, but in areas, such as the Southern Ocean, where they are significant, they may set up quasi-steady ‘lee waves,’
A number of processes can contribute to the filling in of the continuous spectrum typically observed. One seems to be resonant wave–wave interactions: the nonlinear terms, involving u =, in the governing fluid dynamical equation vanish identically for a single wave, but produce interaction terms if two waves are present. These terms, in the momentum and density equations, may be regarded as forcing terms with frequencies and wavenumbers given by the sum and difference frequencies and wavenumbers. For some pairs the sum (or difference) frequency is exactly what would be expected for a free wave with the sum (or difference) wavenumber, so this wave is now resonantly excited, acquiring energy from the original two waves. Detailed calculations for this theory, and using a different approach when the assumptions of weak interaction break down, do not actually make it clear how a typical spectrum arises, but suggest that, once it is present, there is a cascade of energy to waves with shorter vertical scales. As will be discussed later, these shorter waves are more likely to become unstable, break down into turbulence, and cause mixing. The direction in which energy flows in frequency is less clear, though one interaction mechanism, akin to the excitation of a simple pendulum by oscillation of its point of support with twice the natural frequency of the pendulum, can produce small-scale waves with half the frequency of a large-scale parent wave (provided that this half-frequency is still greater than f). Bottom Reflection and Scattering
Internal waves have a frequency less than the local value of the buoyancy frequency N. This typically decreases with increasing depth below the sea surface, so that some downward-propagating waves must undergo internal reflection at a level where their frequency matches the local N. Waves with a frequency less than N at the seafloor (or just above some well-mixed bottom boundary layer) will be scattered and reflected there. The reflection process is unlike that for, say, sound waves, in that for internal waves to conserve their
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INTERNAL WAVES
cgr ki
cgi
i
r kr
Figure 4 Internal wave rays are reflected at the seafloor at an equal angle to the normal to a bottom slope. Subscripts i and r indicate incident and reflected waves.
frequency on reflection, they must, by eqn [1], conserve their angle to the vertical, not their angle to the normal. As a consequence, waves reflected upslope will have a shorter wavelength and narrower ray tube (Figure 4). The latter effect, combined with a reduction in group velocity, causes an increase in wave amplitude, particularly near the ‘critical frequency’ for which the wave rays are parallel to the slope. The waves may be amplified less, or even reduced, for other azimuthal angles of incidence, but it turns out that overall amplification is expected for an isotropic incident spectrum, as has been well documented for internal waves near the steeply sloping sides of Fieberling Guyot in the Pacific Ocean. This analysis certainly applies if the length scale of the slope is large compared with the wavelength. For smaller-scale topography, some energy may be backscattered without as much amplification, but, in general, internal wave interaction with bottom topography tends to redistribute energy toward shorter wavelengths. This may be just as important as wave–wave interactions in shaping the wavenumber part of the internal wave spectrum, though bottom interactions do not affect the frequency distribution. Energetics and Mixing
The general picture that has emerged for internal waves in the ocean is that there is a cascade of energy to smaller scales as a consequence of wave–wave interactions and bottom scattering. This leads to a tendency for shear instability of the horizontal currents, with an expected vertical scale of the order of 1 m for a typical spectrum. It is generally assumed that a fraction of about 15–20% of the energy lost in the breaking leads to an increase in the potential energy of the water column (with the rest of the energy being dissipated and ultimately appearing as a negligible internal heating rate). The associated vertical mixing rate, or ‘eddy diffusivity,’ is of the order
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of 105 m2 s1, again for typical spectral energy levels in the main thermocline. This agrees rather well with estimates based on measurements of turbulent microstructure, and with even more direct estimates based on observations of the vertical spread of an artificial tracer. The mixing may be considerably more intense throughout the water column in regions, such as the Southern Ocean, where internal wave energy levels seem higher, perhaps as a consequence of seafloor generation of the waves by strong mean currents over rough topography. Stronger mixing is also observed in general within a few hundred meters of the bottom in areas of rough bottom topography, though it is not clear whether this results directly from the increased shear of the reflected and scattered waves, or via stronger wave–wave interactions at increased internal wave energies. The relative importance of wind-generated internal waves and internal tides in these regions is also still unsettled, though a topic of active research. Rather weak mixing in the main thermocline of the ocean, together with much weaker stratification in abyssal areas of strong mixing, means that the overall energy loss from the internal wave field is small enough that it would take many tens of days to drain the observed energy levels. This may tie in with the remarkable feature, mentioned earlier, that observed internal wave energy levels in the ocean seem to be rather uniform in space and time, at least much more so than for surface gravity waves; there is no such thing as an ‘internal calm.’ The interpretation is that the decay time of internal waves is, unlike the situation for surface waves, considerably longer than the interval between generation events. There is still some seasonal modulation of the internal wave energy levels, but less than that in the wind and also in accord with a decay time of tens of days. There are exceptions to this picture, of course, with, for example, much lower internal wave energy levels and mixing in the Arctic Ocean (except near some topographic features), perhaps as a consequence of less wind generation, because of the protective ice cover, as well as rather weak tidal currents. The overall dynamical balance of the internal wave field in the ocean is qualitatively summarized in Figure 5. Internal Waves on the Continental Shelf
The above discussion of internal waves has been focused on the deep-sea situation. There are some similarities on the much shallower continental
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Wind
Inertial oscillations
Surface waves
Mixed layer turbulence
Largescale flows
Topography
Continuum
f
Interior mixing
Barotropic tides
M2
Boundary mixing
Figure 5 A summary of internal wave generation, evolution, and eventual dissipation causing mixing in the ocean interior or near boundaries. The dashed lines indicate conjectured energy pathways. (From Mu¨ller and Briscoe (2000)).
shelves, though with a considerable fraction of observed internal wave energy often being associated with the rather large interfacial waves, or their equivalent in a smoothly stratified fluid, discussed at the beginning of this article. It is not clear how much of the rest of the internal wave field on shelves is locally generated and how much propagates there from the deep ocean.
making the shear across the mixed layer base more destablizing during the shallow phase and enhancing the overall mixing. This is an effect that has been excluded from models of the surface layer, and may be a partial reason why these models sometimes need to include ad hoc extra mixing just below the base of the layer.
Conclusions Other Aspects In the atmosphere, internal waves are crucial in establishing the general circulation by transporting momentum from one location to another and then depositing it when they break. One reason for this breaking is that as internal waves propagate vertically into thinner air they must increase their amplitudes in order to conserve their energy flux, and so become more prone to instability. This is not a factor in the oceans, where the density change is very minor. The ratio of mean flow speeds to wave speeds is also less in the ocean than in the atmosphere, making interactions between waves and currents less important in the ocean. None the less, it does seem likely that there are some locations in the ocean where internal wave breaking should drive mean flows. One possible location is the continental slope; internal waves generated as lee waves at one location may propagate shoreward, break on the slope, and drive an along-slope current, much as ocean swell incident at an angle to a beach may drive longshore currents inside the breaker zone. The role of internal waves in other situations may also have been somewhat unrecognized so far. One is their effect on surface mixed layer deepening. The waves alternately shallow and deepen the layer,
Internal waves are both an unavoidable nuisance and a key ingredient of the behavior of the ocean; perhaps they are like the clouds in the atmosphere. The analogy is certainly a good one when one thinks of modeling the large-scale circulation of the two media for applications such as climate prediction. Numerical models fail by several orders of magnitude to have sufficient resolution to treat them explicitly, so their effects must be parameterized. This requires not just an understanding of their role in present conditions, but also a submodel that will predict their characteristics and effects in a changing mean state. A model for internal waves will need to account for the whole awkward mix of generation, propagation, wave–wave interactions, interactions with the mean state, and reflection and scattering from the rough seafloor. We have a partial understanding of many of the pieces but are a long way from putting them all together.
See also Breaking Waves and Near-Surface Turbulence. Internal Tidal Mixing. Internal Tides. Surface Gravity and Capillary Waves. Wave Generation by Wind.
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INTERNAL WAVES
Further Reading Garrett C and Munk WH (1979) Internal waves in the ocean. Annual Review of Fluid Mechanics 11: 339--369. Gill AE (1982) Atmosphere–Ocean Dynamics. New York: Academic Press. Kundu PK (1990) Fluid Dynamics. New York: Academic Press.
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Mu¨ller P and Briscoe M (2000) Diapycnal mixing and internal waves. Oceanography 13: 98--103. Munk WH (1981) Internal waves and small-scale processes. In: Warren BA and Wansch C (eds.) Evolution of Physical Oceanography, pp. 264--291. Princeton: MIT Press.
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INTERNATIONAL ORGANIZATIONS M. R. Reeve, National Science Foundation, Arlington VA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1342–1348, & 2001, Elsevier Ltd.
Introduction This article is limited to a description of the most important organizations related to Ocean Sciences, rather than exhaustively cataloging perhaps hundreds of entities which may cross one or more national borders. I have classified the major international organizations into governmental (or more accurately intergovernmental) and nongovernmental organizations, and within those groupings, global and regional organizations. With the proliferation of the world wide web, much information can be readily accessed from each organization’s web page, some of which I have edited and utilized.
Intergovernmental Organizations (Global) The overarching global intergovernmental organization is the United Nations (UN). Within it, the Intergovernmental Oceanographic Commission (IOC) has the main responsibility for the coastal and deep oceans. Organizationally, the IOC is a component commission of the United Nations Educational, Scientific and Cultural Organization (UNESCO). The World Meteorological Organization (WMO), a UN specialized agency, has strong interactions with the ocean interests, by virtue of the coupling of weather and climate with the circulation of the oceans and its other properties. The Fisheries and Agricultural Organization of the UN also has strong ties to the oceans through its work in marine fisheries.
Intergovernmental Oceanographic Commission (IOC) The work of the IOC, founded in 1960, has focused on promoting marine scientific investigations and related ocean services, with a view to learning more about the nature and resources of the oceans. The IOC focuses on four major themes: (1) facilitation of international oceanographic research programs; (2) establishment and coordination of an operational global ocean observing system; (3) education and
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training programs and technical assistance; and (4) ensuring that ocean data and information are made widely available. The IOC is currently composed of 126 Member States, an Assembly, an Executive Council and a Secretariat. The Secretariat is based in Paris, France. Additionally the IOC has a number of Subsidiary Bodies. Each Member State has one seat in the Assembly, which meets once every two years. The Assembly is the principal organ of the Commission which makes all decisions to accomplish the objectives of the IOC. The Secretariat is the executive arm of the organization. It is headed by an Executive Secretary who is elected by the Assembly and appointed by the Director-General of UNESCO. Countries contribute dues to support the work of the IOC. Scientific/technical subsidiary bodies of IOC include Ocean Science In Relation To Living Resources, Ocean Science In Relation to Non-Living Resources, Ocean Mapping (OM), Marine Pollution Research and Monitoring, Integrated Global Ocean Services System, Global Ocean Observing System, and International Oceanographic Data And Information Exchange. There are various subprograms attached to most of these including the Global Coral Reef Monitoring Network. IOC also has regional subsidiary bodies with responsibilities for carrying out region-specific programs voted by the Assembly. These are the Subcommission for the Caribbean and Adjacent Regions, Regional Committee for the Southern Ocean, Regional Committee for the Western Pacific, Regional Committee for the Cooperative Investigation in the North and Central Western Indian Ocean, Regional Committee for the Central Indian Ocean, Regional Committee for the Central Eastern Atlantic and Regional Committee for the Black Sea. The work of the Secretariat is accomplished by a small permanent staff and a larger number of scientists usually funded by institutions or agencies in their own countries, who have interest in a specific aspect of the activities of the IOC. The funds provided to do this are in addition to country dues. This mechanism to focus staff support on such areas of interest is good, and helps to counterbalance the opposite tendency of the General Assembly, meeting every two years, to direct the General Secretary to undertake new activities. Because such activities must usually be funded out of the existing budget, the tendency inevitably dilutes existing activities, sometimes to a subcritical level.
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INTERNATIONAL ORGANIZATIONS
World Meteorological Organization (WMO) The World Climate Program (WCP) is the main activity of WMO related to the oceans. Established in 1979, the WCP includes the World Climate Research Program (WCRP), of which the Global Climate Observing System, encompassing all components of the climate system, atmosphere, biosphere, cryosphere, and oceans, is a component. WMO and its WCRP are treated fully elsewhere and only a brief summary is provided here, relating to oceans. The World Meteorological Convention, by which the World Meteorological Organization was created, was adopted in 1947, and in 1951 was established as a specialized agency of the United Nations. There at least 185 member countries and territories. The World Meteorological Congress, which is the supreme body of WMO, meets every four years. In order to assess available information on the science, impacts and the cross-cutting economic and other issues related to climate change, in particular possible global warming induced by human activities, WMO and the United Nations Environment Program (UNEP) established the Intergovernmental Panel on Climate Change in 1988. In the early years of their development, WCRP incorporated the Tropical Ocean – Global Atmosphere Study initiated by the Scientific Committee on Oceanic Research (SCOR) (see below), and the World Ocean Circulation Experiment (WOCE). The former was successfully completed and the latter is in its later stages of analysis and synthesis, the major field programs including the WOCE Hydrographic Survey, having been completed. Both programs have been very successful, to the point that an offspring is beginning to be implemented – the Climate Variability and Predictability Study. This 15-year program to observe the atmosphere and oceans will incorporate new technologies developed as a result of the predecessor programs, and rapidly increasing sophistication of computer hardware and software to develop new coupled models. Past climates will also be reconstructed. Related to this effort is the Global Ocean Data Assimilation Experiment which is planned for the first half of this decade. This program was a product of the joint IOC/WCRP Ocean Observations Panel for Climate. In 1999 the governing bodies of WMO and IOC agreed to set up a Joint Technical Commission for Oceanography and Marine Meteorology to act as a coordinating mechanism for the full range of WMO and IOC existing and future operational marine program activities, including the coordinating and
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managing the implementation of an operational Global Ocean Observing System.
The World Bank The Global Environmental Facility (GEF) of the World Bank was launched in 1991 as an experimental facility and evolved ‘to serve the environmental interests of people in all parts of the world’. By 2000, more than 36 nations pledged $2.75 billion in support of GEF’s mission to protect the global environment and promote sustainable development. This has been complemented by $5 billion in co-financing from GEF partners, which include the UN Development Program and the UN Environment Program, as well as host countries. GEF funds projects in four areas: biodiversity, climate change, international waters, and ozone. Up to 1999 the GEF had allocated over $155 million to international waters initiatives. The term ‘international’ refers to fresh as well as ocean waters, and the projects seek to reverse the degradation of bodies of water controlled by a mosaic of regional and international water agreements. A list of currently funded projects is available from the GEF web site. Examples include Black Sea Environmental Management, Gulf of Guinea Large Marine Ecosystem and various coral reef rehabilitation and projection projects. The GEF Secretariat is located within the World Bank in Washington DC, USA.
Intergovernmental Organizations (Regional) Besides regional activities of global intergovernmental organizations referred to above, there are two main regional intergovernmental organizations. The first, the International Council for the Exploration of the Sea (ICES) (for the North Atlantic), and the second is the North Pacific Marine Science Organization. International Council for the Exploration of the Sea (ICES)
Oceanographic investigations form an integral part of the ICES program of multidisciplinary work aimed at understanding the features and dynamics of water masses and their ecological processes. In many instances emphasis is placed on the influence of changes in hydrography (e.g., temperature and salinity) and current flow on the distribution, abundance, and population dynamics of finfish and shellfish stocks. These investigations are also relevant
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to marine pollution studies because physical oceanographic conditions affect the distribution and transport of contaminants in the marine environment. ICES promotes the development and calibration of oceanographic equipment and the maintenance of appropriate standards of quality and comparability of oceanographic data. ICES is the oldest intergovernmental organization in the world concerned with marine and fisheries science. Since its establishment in Copenhagen in 1902, ICES has been a leading scientific forum for the exchange of information and ideas on the sea and its living resources, and for the promotion and coordination of marine research by scientists within its member countries. Each year, ICES holds more than 100 meetings of its various working groups, study groups, workshops, and committees. These activities culminate each September when ICES holds its Annual Science Conference, which attracts 500–1000 government, academic, and other participants. Proceedings of these meetings, and other related activities are published by ICES. Membership has increased from the original eight countries in 1902 to the present 19 countries which come from both sides of the Atlantic and include Canada, the USA and all European coastal states except the Mediterranean countries from Italy eastward. Each country has a vote in the governance, through two delegates, all of whom come together annually at the ‘statutory meeting’ held at the time of the science conference. The Council elects a President at three year intervals from its members, as well as a small executive group (the Bureau) to conduct business intercessionally. The ICES Secretariat in Denmark maintains three databanks, the Oceanographic databank, the Fisheries databank, and the Environmental (marine contaminants) databank. Since the 1970s, a major area of ICES work as an intergovernmental marine science organization has been to provide information and advice to member country governments and international regulatory commissions (including the European Commission) for the protection of the marine environment and for fisheries conservation. This advice is peer-reviewed by the Advisory Committee on Fishery Management and the Advisory Committee on the Marine Environment before being passed on. The structure and operation of ICES has continuously evolved, to meet current needs of the adviceseeking member nations and the oceanographic community. In earlier years it focused mostly on the relationship of oceanography to fisheries, but over the past 15 years, the need has arisen increasingly for advice over the broad range of environmental issues
from marine contaminants to effects of fishing activities on the environment, particularly the seafloor ecosystems. The annual meeting, which was traditionally a business occasion where standing committees reviewed the activities of working groups and passed recommendations on to the Council, has evolved into a major north Atlantic science meeting on oceanography and its application to regional societal problems. ICES has tried to capture its past century of achievements through special lectures, a History Symposium and a soon to be published written history volume. Its future is mapped through the development of its first strategic plan, and ongoing organizational evolution. North Pacific Marine Science Organization (PICES)
PICES held its first Annual Meeting in October 1992, in Victoria, British Columbia. From the beginning, the PICES approach has been multidisciplinary, with standing committees concerned with biological oceanography, fishery science, physical oceanography and climate, and marine environmental quality. There has been growing interaction among these specialties, with joint scientific sessions, interdisciplinary symposia, and a broad study of climate change and carrying capacity (the CCCC program) in the region. Most recently, PICES has taken the lead in joining forces with other international organizations to organize an intersessional Conference under the title of El Nin˜o and Beyond (March 2000). Although PICES is an infant compared with its prototype, ICES, it has already become a major focus for international cooperation in marine science in the northern North Pacific. PICES is an intergovernmental scientific organization. Its present members are Canada, People’s Republic of China, Japan, Republic of Korea, Russian Federation, and the USA. The purposes of the organization are: to promote and coordinate marine research in the northern North Pacific and adjacent seas especially northward of 301N; advance scientific knowledge about the ocean environment, global weather and climate change, living resources and their ecosystems, and the impacts of human activities; and promote the collection and rapid exchange of scientific information on these issues. PICES annual meetings, symposia and workshops provide fora at which marine scientists interested in the North Pacific can exchange latest results, data, and ideas and plan joint research. These meetings have been effective in stimulating and accumulating the interest of the scientific community in member countries to coordinate marine science on the basin
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scale. PICES has been successful at bringing oceanographers from a number of Pacific Rim countries and disciplines (physical, chemical, and biological) together talking and working with fisheries scientists to understand and eventually forecasts temporal variability of ocean ecosystems, and their key species. It has also encouraged and facilitated collaborative marine scientific research in the North Pacific. International collaboration in this region is essential since the open North Pacific is too large for any one country to study adequately on its own, and oceanic circulation and biological species do not recognize international boundaries. Comparisons of similar species and ocean environments on the eastern and western sides of the Pacific will be much more revealing about the large-scale processes affecting ecosystem and fish dynamics than isolated studies in national waters alone. PICES was established, in large measure, by the tireless efforts of Warren Wooster, of the University of Washington, USA. He had long been active in ICES, including as its President, and had become convinced that there was a great need for a similar organization focused on the North Pacific. Other Regional Commissions
Throughout the world, governments have formed regional organizations to protect the environment and regulate activities. Two European examples of these are the Helsinki Commission, otherwise known as the Baltic Marine Environment Protection Commission, and OSPAR Commission for the Protection of the Marine environment of the north-east Atlantic. Their members comprise the countries bordering these specific marine environments. There are also several such commissions organized to protect and regulate specific fisheries and/or fishery regions (e.g., the North Atlantic Salmon Commission), listings and explanations of which go beyond the scope of this article.
Nongovernmental Organizations (Global) International Council for Science (ICSU)
Formerly known as the International Council of Scientific Unions, ICSU is a nongovernmental organization, founded in 1931 to bring together natural scientists in international scientific endeavor. It comprises 95 multidisciplinary National Scientific Members (scientific research councils or science academies) and 25 international, single-discipline Scientific Unions to provide a wide spectrum of
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scientific expertise enabling members to address major international, interdisciplinary issues which none could handle alone. ICSU also has 28 Scientific Associates. The Council seeks to break the barriers of specialization by initiating and coordinating major international interdisciplinary programs and by creating interdisciplinary bodies which undertake activities and research programs of interest to several members. It acts as a focus for the exchange of ideas and information and the development of standards. Hundreds of congresses, symposia and other scientific meetings are organized each year around the world, and a wide range of newsletters, handbooks, and journals is published. The principal source of finance for the ICSU is the contributions it receives from its members. Other sources of income are grants and contracts from UN bodies, foundations, and agencies, which are used solely to support the scientific activities of the ICSU Unions and interdisciplinary bodies. ICSU has a three-tier system of governance. They are the General Assembly (the highest organ), the Executive Board, and the Officers. These are assisted by a Secretariat responsible for the day-to-day work of the Council. Interdisciplinary ICSU bodies are created by the General Assembly as the need for these arises in cooperative projects. Two of these, the International Geosphere-Biosphere Program (IGBP) and SCOR, are currently of particular interest to the ocean sciences community. IGBP, planning for the International GeosphereBiosphere Program: A Study of Global Change, was begun in 1986. The IGBP is a research program with the objective to describe and understand the interactive physical, chemical, and biological processes that regulate the total Earth system, the unique environment that it provides for life, the changes that are occurring in this system, and the manner in which they are influenced by human actions. The program is focused on acquiring basic scientific knowledge about the interactive processes of biology and chemistry of the earth as they relate to global change. Priority is placed on those areas in each of the fields involved that deal with key interactions and significant changes on timescales of decades to centuries, that most affect the biosphere, that are most susceptible to human perturbations, and that will most likely lead to a practical, predictive capability. SCOR, the Scientific Committee on Oceanic Research, established in 1957, is the oldest of ICSU’s interdisciplinary bodies. The recognition that the scientific problems of the oceans required a truly interdisciplinary approach was embodied in plans for
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INTERNATIONAL ORGANIZATIONS
the International Geophysical Year. Accordingly, SCOR’s first major effort was to plan a coordinated, international attack on the least-studied ocean basin of all, the Indian Ocean. The International Indian Ocean Experiment of the early 1960s was the result. For the next 30 years, the reputation of SCOR was largely based on the successes of its scientific working groups. These small international groups of not more than ten members are established in response to proposals from national committees for SCOR, other scientific organizations, or previous working groups. In general, they are designed to address fairly narrowly defined topics (often new, ‘hot’ topics in the field) which can benefit from international attention. Although SCOR does not have the resources to fund research directly, many of its scientific groups have organized international meetings and produced important publications in the scientific literature. Others have proposed and planned large international collaborative efforts such as the Joint Global Ocean Flux Study and Global Ocean Ecosystem Dynamics. Scientists from the 39 SCOR member countries participate in its working groups and steering committees for the larger programs. SCOR often works in association with intergovernmental organizations such as IOC and ICES. The work of SCOR falls into two major categories. The first of these is the traditional mechanism of the SCOR working group, addressing topics which range from the ecology of sea ice to the role of wave breaking on upper ocean dynamics and from coastal modeling to the biogeochemistry of iron in sea water and the responses of coral reefs to global change. For longerterm, complex activities, such as the planning and implementation of large-scale programs SCOR establishes scientific committees. Over the past 15 years there has been a growing awareness of the influence of the ocean in large-scale climate patterns and in moderating global change. By the early 1980s the promise of increased computing capabilities and new satellite instruments for remote sensing of the global ocean permitted oceanographers to conceive of large-scale, internationally planned and implemented experiments of the sort never before possible. The first two of these were the World Ocean Circulation Experiment (WOCE) and the Tropical Ocean–Global Atmosphere Study (TOGA, 1985– 1995). Both grew out of SCOR’s former Committee on Climatic Changes and the Ocean which was also cosponsored by the IOC. A few years ago WOCE was incorporated into the World Climate Research program; its field program is now complete and WOCE is now embarking upon the critical phase of analysis, interpretation, modeling and synthesis.
Since the late 1980s SCOR has played a major role in fostering the development of two newer global change programs, both of which now form part of the IGBP effort. These are the Joint Global Ocean Flux Study and Global Ocean Ecosystem Dynamics. SCOR consists of its ‘members’ – the national committees for oceanic research of its 39 member countries, each of which is represented by three individual oceanographers. The biennial general meetings elect an executive committee, which also includes ex officio members from allied disciplinary organizations, namely, the International Association for Physical Sciences of the Ocean, the International Association for Biological Oceanography, and the International Association for Meteorological and Atmospheric Sciences. The Ocean Drilling Program (ODP)
The ODP is an international partnership of scientists and research institutions organized to explore the evolution and structure of the earth. It uses drilling and data from drill holes to improve fundamental understanding of the role of physical, chemical, and biological processes in the geological history, structure, and evolution of the oceanic portion of the earth’s crust. The ODP provides researchers around the world access to a vast repository of geological and environmental information recorded far below the ocean surface in seafloor sediments and rocks. The National Science Foundation (US Federal Government) supports approximately 60% of the total international effort. Other partners (Germany, France, Japan, the UK, the Australia/Canada/Chinese Taipei/Korea consortium, European Science Foundation consortium, and the People’s Republic of China), comprising 20 other countries, provide 40% of the program costs. Currently, specific studies include documenting the history of volcanic plumes in the western Pacific, examining the formation of mineral deposits near west Pacific island arcs, instrumentation of boreholes to study seismicity of the north-west Pacific, and recovery of gas hydrate deposits to examine their formation along the Oregon margin. Support will also be provided for new scientific and operational developments to extend capabilities for deep biosphere investigations for ocean biocomplexity studies. Other Nongovernmental Organizations (Global and Regional)
Within this category there are many scientific unions and societies, as well as other organizations, which are either explicitly ocean-oriented or have ocean
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INTERNATIONAL ORGANIZATIONS
components. Several, as noted above are related to ICSU. Others include the American Society of Limnology and Oceanography and the American Geophysical Union (both primarily north American, but with a substantial international membership), the European Geophysical Society, the Oceanography and Challenger Societies, and many others.
•
Acknowledgements
•
I wish to gratefully acknowledge the assistance of my colleague Kandace Binkley in the preparation of this article.
•
Memorial
•
I dedicate this article as a small memorial for my friend and colleague Dr. George Grice, whose untimely death occurred in March 2001. A former Associate Director of the Wood’s Hole Oceanographic Institution and Deputy Director of the Northeast Fisheries Science Center in Wood’s Hole, George was involved with international organizations all his professional life, particularly ICES and IOC. He was serving the latter institution at the time of his death as Senior Science Advisor to the Executive Secretary, Dr. Patricio Bernal.
• • •
•
• • • • • •
Links to International Organizations
• • • • • •
American Society of Limnology and Oceanography (ASLO) – http://aslo.org American Geophysical Union (AGU) – http:// www.agu.org Challenger Society for Marine Science – http:// www.soc.soton.ac.uk/OTHERS/CSMS European Geophysical Society (EGS) – http:// www.mpae.gwdg.de/EGS/EGS.html Global Ocean Ecosystem Dynamics (GLOBEC) – http://www1.npm.ac.uk/globec International Association for Meteorology and Atmospheric Sciences (IAMAS) – http:// iamas.org
• • • • • •
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International Association for the Physical Sciences of the Oceans (IAPSO) – http://www. olympus.net/IAPSO International Council for the Exploration of the Sea (ICES) – http://www.ices.dk International Council for Science (ICSU) – http:// www.icsu.org International Geosphere-Biosphere Program (IGBP) – http://www.igbp.kva.se Intergovernmental Oceanographic Commission (IOC) – http://ioc.unesco.org/iocweb Joint Global Ocean Flux (JGOFS) – http:// ads.smr.uib.no/jgofs/jgofs.htm Baltic Marine Environment Protection Commission (HELCOM) – http://www.helcom.fi/ oldhc.html North Pacific Marine Science Organization (PICES) – http://pices.ios.bc.ca Ocean Drilling Program (ODP) – http://wwwodp.tamu.edu OSPAR Commission for the Protection of the Marine Environment of the North-East Atlantic – http://www.ospar.org Scientific Committee on Antarctic Research (SCAR) – http://www.scar.org Scientific Committee on Oceanic Research (SCOR) – http://www.jhu.edu/Bscor The Oceanography Society (TOS) - http://tos.org Tropical Ocean – Global Atmosphere Coupled Ocean/Atmosphere Response Experiment (TOGA) – http://trmm.gsfc.nasa.gov/trmm_office/field_campaigns/toga_coare/toga_coare.html United Nations (UN) – http://www.un.org United Nations Environment Program (UNEP) – http://www.unep.ch/index.html United Nations Educational, Scientific and Cultural Organization (UNESCO) – http:// www.unesco.org World Bank, Global Environmental Facility (GEF) – http://www.gefweb.org World Meteorological Organization (WMO) – http://www.wmo.ch World Ocean Circulation Experiment (WOCE) – http://www.soc.soton.ac.uk/OTHERS/woceipo/ ipo.html
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INTERTIDAL FISHES R. N. Gibson, Scottish Association for Marine Science, Argyll, Scotland Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1348–1354, & 2001 Elsevier Ltd.
Introduction and Classification of Intertidal Fishes The intertidal zone is the most temporally and spatially variable of all marine habitats. It ranges from sand and mud flats to rocky reefs and allows the development of a wide variety of plant and animal communities. The members of these communities are subject to the many and frequent changes imposed by wave action and the ebb and flow of the tide. Consequently, animals living permanently in the intertidal zone have evolved a variety of anatomical, physiological and behavioral adaptations that enable them to survive in this challenging habitat. The greater motility of fishes compared with most other intertidal animals allows them greater flexibility in combating these stresses and they adopt one of two basic strategies. The first is to remain in the zone at low tide. This strategy used by the ‘residents’, requires the availability of some form of shelter to alleviate the dangers of exposure to air and to predators. ‘Visitors’ or ‘transients’, that is species not adapted to cope with large changes in environmental conditions, only enter the intertidal zone when it is submerged and leave as the tide ebbs. The extent to which particular species employ either of these strategies varies widely. Many species found in the intertidal zone spend most of their lives there and are integral parts of the intertidal ecosystem. At the other extreme, others simply use the intertidal zone at high tide as an extension of their normal subtidal living space. In between these extremes are species that spend seasons of the year or parts of their life history in the intertidal zone and use it principally as a nursery or spawning ground. The different behavior patterns used by residents and visitors mean that few fishes are accidentally stranded by the outgoing tide.
Habitats, Abundance, and Systematics Fishes can be found in almost all intertidal habitats and in all nonpolar regions. Most shelter is found on rocky shores in the form of weed, pools, crevices and
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spaces beneath boulders and it is on rocky shores that resident intertidal fishes are usually most numerous. If, however, fishes are capable of constructing their own shelter in the form of burrows in the sediment, as in the tropical mudskippers (Gobiidae), they may be abundant in such habitats. Visiting species may also be extremely numerous on occasions and are particularly common in habitats such as sandy beaches, mudflats and saltmarshes. Few fishes occupy gravel beaches although species like the Pacific herring (Clupea pallasii), capelin (Mallotus villosus, Osmeridae) and some pufferfishes (Tetraodontidae) may spawn on such beaches. Estimating abundance in terms of numbers per unit area can be difficult because of the cryptic nature of the fishes and the patchiness of the habitat. The difficulty is particularly acute on rocky shores where fishes may be highly concentrated in areas such as rock pools but absent elsewhere. Nevertheless, fish densities can be relatively high, particularly at the time of recruitment from the plankton, and on occasions may exceed 10 individuals per m2. Over 700 species of fishes from 110 families have so far been recorded in the intertidal zone worldwide. This figure represents less than 3% of known fish species but is certainly an underestimate because it is based only on species recorded on rocky shores. Species on soft sediment shores are not included and many areas of the world have yet to be studied in detail. The final count is therefore likely to be much higher. Intertidal fish faunas are frequently dominated by members of a few families (Table 1). Generally speaking, more species are found in the tropics than in temperate zones and each area of the world tends to have its own characteristic fauna. The Atlantic coast of South Africa, for example, is characterized by large numbers of clinid species, New Zealand by triplefins, the northeast Pacific by sculpins and pricklebacks (Stichaeidae) and many other areas by blennies, gobies and clingfishes.
Characteristics of Intertidal Fishes as Adaptations to Intertidal Life Resident intertidal fishes are probably descended from subtidal ancestors and have few, if any, characters that are truly unique. They are thus representatives of families that have convergently evolved morphological, behavioral and physiological traits that enable them to survive in shallow turbulent
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Table 1 Analysis of 47 worldwide collections of rocky shore intertidal fishes to show the 10 families with the largest number of species. Based on Prochazka et al. in Horn et al. (1999). Note that abundance of species does not necessarily imply that families are also numerically abundant
mudskippers of the genus Periophthalmus and blennies of the genera Alticus and Coryphoblennius.
Family
Common name
Number of species
Blenniidae Gobiidae Labridae Clinidae Pomacentridae Tripterygiidae Cottidae Labrisomidae Scorpaenidae Gobiesocidae
Blennies and rockskippers Gobies and mudskippers Wrasses Clinids, kelpfishes, klipfishes Damselfishes Triplefin blennies Sculpins Labrisomids Scorpionfishes Clingfishes
55 54 44 33 30 30 26 26 25 24
All the common families of intertidal fishes possess the characteristic morphological features of fishes adapted for benthic life in turbulent waters. They are cryptically colored, are rarely more than 15 cm long and are negatively buoyant because they lack a swimbladder or possess one that is reduced in volume. Four basic body shapes can be recognized: elongate, dorsoventrally flattened (depressed), smoothly cylindrical (terete), or laterally compressed (Figure 1). In many species the fins are modified to act as attachment devices to prevent dislodgement by turbulence or to assist movement over rough surfaces. In most blennies, the rays of the paired and ventral fins are hooked at their distal ends and may be covered with a thick cuticle to minimize wear. In the gobies and clingfishes, the pelvic fins are fused to form suction cups and allow the fish to attach themselves firmly to the substratum. There seem to be few ‘typical’ sensory adaptations to intertidal life although many species have reduced olfactory and lateral line systems. These sensory systems would be of limited use for species living in turbulent waters or in those that frequently emerge from the water.
habitats. The distribution of many resident intertidal species also extends below low water mark but they mainly differ from their fully subtidal relatives in the degree to which they can withstand exposure to air and are capable of terrestrial locomotion. Nevertheless, a few species can be considered truly intertidal in their distribution because they never occur below low water mark. Examples are the
Morphology
Behavior
(A)
(B)
(C)
(D)
Figure 1 Sketches of four intertidal fish species to demonstrate the basic body shapes. (A) Terete (Gobius paganellus, Gobiidae); (B) dorsoventrally flattened (Lepadogaster lepadogaster, Gobiesocidae); (C) elongate (Pholis gunnellus, Pholidae); (D) laterally compressed (Symphodus melops, Labridae).
Intertidal fishes also show characteristic behavior that enables them to cope with the rigors of intertidal life. Their modified fins and relatively high density allow them to remain on or close to the bottom with the minimum of effort and to resist displacement by surge. Most are also thigmotactic, a behavior that keeps as much of their body touching the substratum as possible and ensures that they come to rest in contact with solid objects when inactive. Their mode of locomotion also reflects this bottom-dwelling lifestyle. Few excursions are made into open water and those species with large pectoral fins use them as much as the tail for forward movement and swim in a series of short hops. Clingfishes and gobies can progress slowly over horizontal and vertical surfaces using their sucker. Elongate forms, which usually have reduced paired fins, creep along the bottom using sinuous movements of their body or alternate lateral flexions of the tail. Strong lateral flips of the tail are also used by some blennies and gobies that can jump between rock pools at low tide and by the semiterrestrial mudskippers in their characteristic ‘skipping’ movements over the surface of the mud. When progressing more slowly mudskippers use the
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muscular pectoral fins as ‘crutches’ to move over the substratum. Resident species are generally active at a particular state of the tide, although these tidally phased movements can be modulated by the day/night cycle so that some species may only be active, for example, on high tides that occur during the night. Visitors present a more complicated picture because, although their movements are also basically of tidal frequency, they are modulated by a wider range of cycles of lower frequency. Individuals may migrate intertidally on each tide, on every other tide depending on whether they are diurnal or nocturnal, or only on day or night spring tides. They may enter the zone as juveniles in spring or summer, stay there for several months, during which time they are tidally active, and then leave when conditions become unsuitable in winter or as they grow and mature. The distances over which fishes move in the course of their intertidal movements are dependent on several factors. At one end of the scale are relatively gradual but ultimately extensive shifts in position related to the seasons. At the other end are local, short-term, tidally related foraging excursions. Residents tend to be very restricted in their movements and are often territorial, whereas visitors regularly enter and leave the intertidal zone and may cover considerable distances in each tide. Body size also determines the scale of movement. Residents are small and have limited powers of locomotion whereas visitors usually possess good locomotory abilities and can travel greater distances more rapidly. The small size and poor swimming abilities of resident species partially account for the restricted extent of their movements but there is good evidence that some species also possess good homing abilities. Most evidence comes from experiments in which individuals are experimentally displaced short distances from their ‘home’ pools and subsequently reappear in these pools a short time later. Experiments with the goby Bathygobius soporator suggest that this species acquires a knowledge of its surroundings by swimming over them at high tide. It can remember this knowledge for several weeks and use its knowledge to return to its pool of origin. Homing is also known in some species of blennies and sculpins and is presumably based on the use of visual clues in the environment. Displacement experiments with the sculpin Oligocottus maculosus, however, suggest that this species at least may also use olfactory clues to find its way back to its home pool. The energy expended in these movements at the various temporal and spatial scales described suggest that they play an important part in the ecology of
both resident and visiting species alike. Several functions have been proposed for these movement patterns of which the most obvious is feeding. Visitors move into the intertidal zone on the rising tide to take advantage of the food resources that are only accessible at high water and move out again as the tide ebbs. Residents, on the other hand, simply move out of their low-tide refuge, forage while the tide is high, and return to the refuge before low tide. Following the flooding tide into the intertidal zone may have the added benefit of providing protection from larger predators in deeper water. Movements at both short and long time scales may also be in response to changing environmental conditions. Visitors avoid being stranded above low water mark at low tide because of the lack of refuges or because they are not adapted for the low tide conditions that may arise in possible refuges such as rock pools. Longer term seasonal movements into deeper water can be viewed as responses to changes in such physical factors as temperature, salinity and turbulence. Finally, several species whose distribution is basically subtidal move into the intertidal zone to spawn (see Life histories and reproduction below). In order to synchronize their behavior with the constantly changing environment fishes must be able to detect and respond to the cues produced by these changes. The cues that fish actually use in timing and directing their tidally synchronized movements are mostly unknown. Synchronization could be achieved by a direct response to change. The flooding of a tide pool or the changing pressure associated with the rising tide, for example, could be used to signal the start of activity. In addition, behavior may be synchronized with the external environment by reference to an internal timing mechanism. The possession of such a ‘biological clock’ that is phased with, but operates independently of, external conditions is a feature common to all intertidal fishes in which it has been investigated. In the laboratory the presence of the ‘clock’ can be demonstrated by recording the activity of fish in the absence of external cues. Under these conditions fish show a rhythm of swimming activity in which periods of activity alternate with periods of rest (Figure 2). In most cases the period of greatest activity appears at the time of predicted high tide on the shore from which the fish originated. These ‘circatidal’ activity rhythms, so called because in constant conditions the period of the rhythm only approximates the average period of the natural tidal cycle (12.4 h), can persist in the laboratory for several days without reinforcement by external cues. After this time activity becomes random but in the blenny Lipophrys pholis the rhythm can be restarted (entrained) by exposing fish to
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80
Activity per hour
Cebidichthys violaceus (Monkeyface prickleback)
60 Large fish
Periophthalmus cantonensis (Chinese mudskipper)
40 20
Small fish
Tomicodon humeralis (Sonora clingfish)
0 0
12
24
36 Time (h)
48
60
72
Figure 2 The ‘circatidal’ activity pattern shown by the blenny Lipohrys pholis in constant laboratory conditions. Peaks of activity correspond initially to the predicted times of high tide (dotted lines) but gradually occur later because the ‘biological clock’ has a period greater than that of the natural tidal cycle (12.4 h). In the sea the clock would be continually synchronized by local tidal conditions.
experimental cycles of wave action or hydrostatic pressure or by replacing them in the sea. The probable function of this circatidal ‘clock’ is that it enables a fish to anticipate future changes in tidal state and regulate its activity and physiology accordingly. If activity is prevented in the wild, by storms for example, then the persistent nature of the clock allows fish to resume activity that is appropriate to the tidal state at the next opportunity. Physiology
The ebbing tide exposes the intertidal zone to air and so the location of a fish is critical to its survival. Over the low tide period resident fish face not only the danger of exposure to air but also to marked changes in other physical and chemical conditions. On rocky shores, pools act as low-tide refuges but even here temperature, salinity, pH and oxygen content of the water can change markedly for the few hours that the pool is isolated from the sea. Consequently, many resident species are usually more tolerant of changes in these factors than subtidal species. Exposure to air could result in desiccation but, surprisingly, resident intertidal fishes show no major physiological or anatomical adaptations for resisting desiccation. Instead, desiccation is minimized by behavior patterns that ensure fish hide in pools, or in wet areas under stones and clumps of weed at low tide. Nevertheless, many species can survive out of water in moist conditions for many hours and tolerate water losses of more than 20% of their body weight, equivalent to that of some amphibians (Figure 3). The ability to tolerate water loss is generally correlated with position on the shore; those fish that occupy higher
Pherallodiscus funebris (Fraildisc clingfish)
Rana clamitans (Green frog)
Scaphiopus couchi (Spadefoot toad)
0
30 10 20 40 % wet body weight
50
Figure 3 Tolerance of water loss as a percentage of wet body weight in four intertidal fishes compared with two amphibians. (Reproduced with permission from Horn and Gibson, 1988.)
levels are the most resistant. Prolonged emersion could also present fish with problems of nitrogen excretion and osmoregulation but those species that have been investigated seem to be able to cope with any changes in their internal medium caused by the absence of water surrounding them. A further consequence of emersion is the change in the availability of oxygen. Although air contains a greater percentage of oxygen than water its density is much lower causing the gill filaments to collapse and reducing the area of the primary respiratory surface. Unlike some freshwater fishes, intertidal fishes that leave the water or inhabit regions where the water is likely to become hypoxic have no specialized airbreathing organs but maximize aerial gas exchange in other ways. Some species have reduced secondary gill lamellae, thickened gill epithelia and their gills are stiffened with cartilaginous rods. Such features reduce the likelihood of gill collapse when the fish is out of water. The skin is also used as an efficient respiratory surface because it is in contact with air and is close to surface blood vessels. In order to be effective, however, the skin must be kept moist and fish that are active out of water frequently roll on their sides or return to the sea to wet the skin. It is probable that some species use vascularized linings of the mouth, opercular cavities, and, possibly, the esophagus as respiratory surfaces. The ability to respire in air is currently known for at least 60 species from 12 families and many of these voluntarily leave the water. Fish capable of
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respiring out of water have been classified into three main types. ‘Skippers’ are commonly seen out of water and actively feed, display and defend territories on land. They are typified by the tropical and subtropical mudskippers (Gobiidae) and rockskippers (Blenniidae). The second much less terrestrially active group, the ‘tidepool emergers’, crawl or jump out of tide pools mainly in response to hypoxia. Species from several families show this behavior but it has been best studied in the sculpins. The third group, the ‘remainers’, comprises many species that sit out the low-tide period emersed beneath rocks and in crevices or weed clumps. Some may be guarding egg masses and they are not active out of water unless disturbed.
Feeding Ecology and Predation Impact Intertidal fishes are no more specialized or generalized in their diets and feeding ecology than subtidal fishes. Most are carnivorous or omnivorous and a few are herbivores. Herbivores appear to be less common in higher latitudes but a satisfactory explanation of this phenomenon is still awaited. In some cases the diet changes with size so that the youngest stages are carnivorous but larger amounts of algae are included in the diet as the fish grow. The small size of most resident fishes also means that their diet is composed of small items such as copepods and amphipods. In common with many other small fishes, they may also feed on parts of larger animals such as the cirri of barnacles or the siphons of bivalve molluscs, a form of browsing that does not destroy the prey. The extent to which intertidal fishes have an impact on the abundance of their prey and on the structure of intertidal communities is not clear. In some areas no impact has been detected whereas in others fish may have a marked effect, particularly on the size and species composition of intertidal algae. Those species that only enter the intertidal zone to feed contribute to the export of energy from this area into deeper water.
Life Histories and Reproduction The majority of resident intertidal fishes rarely live longer than two to three years although some temperate gobies and blennies have a maximum life span of up to 10 years. Maturity is achieved in the first or second year of life and the females of longer-lived species may spawn several times a year for each year thereafter. Representatives of 25 teleost families are known to spawn in the intertidal zone. Of these,
residents and visitors make up about equal proportions. Intertidal spawning has both costs and benefits. For resident species, intertidal spawning reduces the likelihood of dispersal of the offspring from the adult habitat. It also obviates the need for movement to distant spawning grounds; a process that would be energetically costly for small-bodied demersal fishes with limited powers of locomotion and would at the same time expose them to greater risks of predation. These advantages do not apply to subtidal species many of which are good swimmers and whose adult habitat is offshore. For these species, the benefits are considered to be reduced egg predation rates and possibly faster development if the eggs become emersed. In both residents and visitors alike eggs spawned intertidally may be subject to the costs of increased mortality caused by desiccation and temperature stress. In addition, visitors may be vulnerable to avian and terrestrial predators during the spawning process. The rugose topography of rocky shores and the cryptic sites chosen for spawning has led to the development of complex mating behavior in many species, particularly the blennies and gobies. In these species, courtship displays by the male include elements of mate attraction and a demonstration of the location of the chosen spawning site. Observing these displays is difficult in most groups but field observations have been made on several Mediterranean blennies that live in holes in rock walls and on mudskippers that perform their mating behavior out of water on the surface of the mud. All species spawn relatively few (range approximately 102–105) large eggs (B 1 mm diameter) that are laid on or buried in the substratum. Large eggs produce large larvae which may reduce dispersal from shallow water by minimizing the amount of time spent in the planktonic stage. On hard substrata the eggs are laid under stones, in holes and crevices and in or under weed. They may be attached individually to the substratum surface in a single layer (blennies, gobies, clingfishes) or in a clump (sculpins). In the gunnels and pricklebacks the eggs adhere to each other in balls but not to the surface. In soft sediments the eggs may be buried by the female as in the grunions (Atherinidae), or laid in burrows as in the mudskippers and some gobies. Killifishes (Fundulus) lay their eggs in salt-marsh vegetation. In temperate latitudes spawning usually takes place in the spring and early summer, but spawning rhythms of shorter frequency may be superimposed on this annual seasonality. Subtidal species such as the grunions, capelin and pufferfishes that use the intertidal zone as a spawning ground mostly take advantage of spring tides to deposit their eggs in the
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INTERTIDAL FISHES
90
Mean % of catch at each distance
sediment high on the shore, usually at night. During the reproductive season such fish, therefore, spawn at fortnightly intervals at the times of the new and full moons. The larvae develop over the intervening weeks and hatch when the eggs are next immersed. Buried eggs are left unattended but eggs laid in layers, clumps or balls are always cared for by the parent. The sex of the individual that undertakes this parental care varies between species. In some it is the male, in others the female and in yet others both sexes participate. In some members of the families Embiotocidae, Clinidae and Zoarcidae fertilization is internal and the young are produced live. Parental care by oviparous species takes a variety of forms but all have the function of increasing egg survival rates and removal of the guardian parent greatly increases mortality. Most species guard the eggs against predators but some also clean and fan the eggs to maintain a good supply of oxygen and reduce the possibility of attack by pathogens. Development time of the larvae depends on species and temperature but when fully formed the eggs hatch to release free-swimming planktonic larvae. In only one case, the plainfin midshipman (Porichthys notatus, Batrachoididae) does the male parent also care for the larvae, which in this species remain attached to the substratum near the nest site. The factors stimulating hatching are mostly unknown although it has been suggested that wave shock and temperature change associated with the rising tide may be involved. Until recently it was assumed that the hatched larvae were dispersed randomly by currents and turbulence. It has been shown, however, that at least some species minimize this dispersion by forming schools close to the bottom and are rarely found offshore (Figure 4). On completion of the larval phase the larvae metamorphose into the benthic juvenile phase and settle on the bottom. The clues used by settling larvae to select the appropriate substratum are poorly known but there is some evidence to suggest that individuals can discriminate
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80 70
Intertidal families Offshore families
60 50 40 30 20 10 0 0
4 20 Distance from shoreline (m)
500
Figure 4 Comparison of the distribution of fish larvae from four intertidal and five offshore families caught in plankton nets at four distances from the shoreline in Vancouver harbour, British Columbia. (Based on data in Marliave JB (1986) Transactions of the American Fisheries Society 115: 149–154.)
between substratum types and settle on their preferred type.
See also Fish Ecophysiology. Fish Larvae. Mangroves. Rocky Shores. Salt Marshes and Mud Flats. Sandy Beaches, Biology of.
Further Reading Gibson RN (1996) Intertidal fishes: life in a fluctuating environment. In: Pitcher TJ (ed.) The Behaviour of Teleost Fishes, 2nd edn, pp. 513--586. London: Chapman Hall. Horn MH and Gibson RN (1988) Intertidal fishes. Scientific American 256: 64--70. Horn MH, Martin KLM, and Chotkowski MA (eds.) (1999) Intertidal Fishes: Life in Two Worlds. San Diego: Academic Press.
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INTRA-AMERICAS SEA latitude, and l ¼ 50–981W longitude. Figure 1 summarizes the geographical setting. Early oceanographic explorations of the region were by European scientists who chose to name the IAS (the Caribbean Sea in particular) the ‘American Mediterranean’. While superficially this terminology describes the IAS as a similar semi-enclosed sea where evaporation (E) exceeds precipitation (P) plus river runoff (R), E > P þ R, the Mediterranean Sea is markedly different in character from its western Atlantic counterpart. Also, the IAS was broken into smaller components, and little attention was paid to the Caribbean Sea and by Gulf of Mexico oceanographers and vice versa. Conversely, the Straits of Florida and the water currents of the Gulf Stream system are perhaps the most widely studied oceanographic features on Earth. With the coming of significant international cooperation between scientists from throughout all the Americas, the IAS began to be appreciated as a
G. A. Maul, Florida Institute of Technology, Melbourne, FL, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1354–1363, & 2001, Elsevier Ltd.
Introduction The Intra-Americas Sea (IAS) is a semi-enclosed saltwater body of the tropical and subtropical western Atlantic Ocean that comprises the Caribbean Sea, the Gulf of Mexico, the Straits of Florida, the Bahamas, the Guianas, and the adjacent waters. Biogeographically, the IAS includes the estuarine, coastal, shelf, and pelagic waters from the mouth of the Amazon River at the equator off Brazil, to Bermuda and westward to the shores of North, Central, and South America. Geographically, the boundaries may be set approximately as f ¼ 01 to 321N
35˚N USA Charleston BERMUDA
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Figure 1 The Intra-Americas Sea, a semi-enclosed water body of the subtropical and tropical western North Atlantic Ocean. Place names in accord with the US Board on Geographic Names.
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INTRA-AMERICAS SEA
unified body of water distinctly different from the early and mid-twentieth century European perspective. An inclusive term was needed to integrate not only the oceanography of the region, but its meteorology and maritime socioeconomic connectivity as well. Thus the term Intra-Americas Sea was developed from the multilingual and multicultural heritage shared by its people.
Regional Overview Pre-Columbian indigenous peoples of the IntraAmericas Sea certainly knew of the oceanic currents and atmospheric winds. Caribes from the south, Tainos in the middle Antilles, Arawaks from the north, Mayas and Olmecs to the west, all moved freely from island to island to continent, presumably with some knowledge of the currents and winds we have named Caribbean Current, Trade Winds, Gulf Stream, Hurricane, and Guianas Current. European explorers and conquistadors relearned this information, not from the IAS’s inhabitants, but from hardship after experiencing what was so well known already. James A. Mitchner’s 1989 novel Caribbean imagines so well what science could have learned directly. As regards the geological setting, the IAS encompasses three tectonic plates: the North American Plate, the Caribbean Plate, and the South American Plate (the Cocos and Nazca Plates mark Pacific tectonic boundaries but are not significantly involved in the air–sea regime discussed herein). About 3 Ma the Caribbean Plate drifting from west to east closed the gap between North and South America, creating Panama and deflecting oceanic flow northward. Central America and the eastern Caribbean margin are volcanically active today. Associated with tectonics are earthquakes and seismic sea waves (tsunami) that are part of the circulation regime, as well as the complex bottom topography that channels water movement and perturbs the atmospheric circulation. Geological forces then not only form the coasts and islands of the IAS; they are a central element in appreciating the flow of air and water. Geophysical fluids obey P P Newton’s laws of motion, specifically F ¼ m a (where F is force, m is mass, and a is acceleration), the laws of thermodynamics, and continuity of mass. The air and water of the IAS are accordingly connected to all of Earth’s ocean–atmosphere continuum, and have special complexities as they flow around and through the passages, channels, and straits within the region. Thus detailed knowledge of the water depths and land heights, and their attendant frictional characteristics, is essential to appreciating the flow patterns discussed below.
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Sill depths control much of the oceanic flow patterns in the IAS. It has been reported that the deepest sill is in the Yucatan Channel (2040 m) between Mexico and Cuba, that the average sill-depth of the Antillean Arc is 1200 m, but that the Jungfern-Anegada Passage is 1815 m deep (between the Virgin Islands and the Lesser Antilles), and that the Windward Passage separating Cuba and Hispaniola is 1690 m deep. Within the IAS are many much deeper basins than these controlling sills, leading to the notion that the waters deeper than the sill depths are moved by convection rather than advection. This infers that the IAS has a two-flow regime: an upper layer mostly influenced by wind-driven advection, and a deep layer controlled by overflows. The controlling sill depth of the Straits of Florida is about 800 m, which means that the outflow of the IAS is topographically accelerated. Figure 2 summarizes the IAS water depth information. Lastly, to appreciate the ocean currents of the IntraAmericas Sea, the structure of atmospheric forcing needs to be mentioned. The southern portion of the IAS is under the influence of the Inter-Tropical Convergence Zone, the ITCZ. In the Northern Hemisphere summer, the ITCZ migrates northward and the easterly Trade Winds at the southern boundary of the Bermuda meteorological high-pressure zone dominate the IAS to its northern extent. As boreal winter approaches, the ITCZ migrates southward and the midlatitude frontal passages sweep across the IAS to south of Cuba and sometimes almost to South America. During early autumn, the atmosphere is characterized by a series of tropical cyclones, that from time to time reach intensities known as the West Indian Hurricane. These severe atmospheric disturbances are cyclonic circulation features that bring not only strong winds, storm surge, and the associated damage, but also much needed precipitation and flushing of shallow bays and estuaries. Climatologically, the Ko¨ppen classification system would place the northern IAS in a Cfa (humid subtropical) category; coastal Central America, the Greater Antilles and the Bahamas as Aw (tropical wet and dry) with sections as Af (tropical rain forest) including the Lesser Antilles. The southern and southeastern coastal IAS is classified as BSh and BW (semiarid or steppe, and arid desert, respectively), with Am (tropical monsoon) along the coasts of the Guianas and Brazil. These classifications are perturbed by interannual and decadal climate oscillations, in particular by El Nin˜o–Southern Oscillation (ENSO) events, which tend to cause cold wet winters in the northern IAS (particularly Florida and Cuba) and warm dry autumn conditions in Panama, coastal Colombia and Venezuela, the Guianas, and northern Brazil.
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30˚N
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Figure 2 Bottom topography of the Intra-Americas Sea; water depths in meters.
Surface Throughflow Regime Perhaps the best way to envision IAS surface currents is to take a Lagrangian perspective, and imagine floating on a northbound satellite-tracked buoy passing the equator off Brazil as part of the ‘oceanic conveyor belt’. As the water parcel travels northward, perhaps entraining some Amazon River water, the dominant physics is conservation of potential vorticity, d=dt½B þ f =H ¼ 0, where relative vorticity B ¼ @v=@x @u=@y, the Coriolis parameter f ¼ 2O sin f with latitude f, and water depth H. For a given H, as the parcel flows northward, f increases and z must decrease, forcing an anticyclonic (clockwise) turning. The region where this occurs is called the North Brazil Current ‘retroflection’ and can be seen in satellite images as a distinct offshore turning of the current. Some of this water continues toward the east in the North Atlantic Equatorial CounterCurrent, but some advects up the South American coast in the Guianas Current to the Lesser Antilles Arc, sometimes as an anticyclonic eddy. Much of this is evident in Figure 3. The northward-flowing water parcel usually passes Barbados and often flows through the passages of
the Lesser Antilles, carrying anticyclonic vorticity and perhaps Amazon riverine particles and biota into the Caribbean Sea. Under the influence of the Northeast Trade Winds, the Caribbean Current moves westward at a leisurely pace of perhaps 0.2 m s1, but with notable meandering and eddying along the path. Sometimes under this same wind regime, water from the Orinoco River is seen to be carried completely across the eastern Caribbean Sea to Puerto Rico and Hispaniola, particularly in late summer. Similarly, in the Panama-Colombia Bight, the wind-stress (t) curl, @ty =@x @tx =@y, causes an rE150 km radius eddy to spin-up and spin-down annually, the so-called Panama-Colombia Gyre (PCG). In the vicinity of the PCG a major South American river, Colombia’s Magdalena, flows into the ocean where its waters and its flotsam mix with the sea, and are carried to distant shores by ocean currents. The PCG is but one feature of the IAS surface current variability now being simulated in numerical models, and being observed by systems such as satellite altimeters and radiometers. As the meandering, eddying, Caribbean Current approaches the Central America coast, it is forced anticyclonically northward into the Yucatan
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Figure 3 Ocean color composite of the Intra-Americas Sea showing concentration of chlorophyll þ phaeophytins and suspended sediments. Image from observations of the Coastal Zone Color Scanner (NASA) compiled by Frank Muller-Karger for October 1979.
Channel. In the area off Belize and Yucatan Mexico, this IAS current takes on the characteristics of the Gulf Stream: a deep (zE1200 m) western boundary current with swift surface flows of more than n ¼ 1 m s1 and a distinct cyclonic horizontal velocity shear boundary, @ n=@x, along the western edge. Here the stream is known as the Yucatan Current, and it is a northward flow connecting the Caribbean Sea with the eastern Gulf of Mexico. Gallegos (1996) has shown that the Yucatan Current is highly geostrophic, the balance of forces (per unit mass) being f v ¼ 1=r @p=@ n, where r is the density of sea water, and n is the current speed at right angles to the horizontal pressure gradient, @p=@ n . Gallegos also studied the temperature–salinity (T–S) structure of the Yucatan Current and has concluded that it has T–S properties similar to those in the offing of Cape Hatteras. As this branch of the Gulf Stream system flows northward, it forces a vigorous upwelling regime along the eastern Campeche Bank that supports one of the IAS’s greatest fisheries. Once in the Gulf of Mexico, the Yucatan Current is known as the Gulf Loop Current because of its characteristic anticyclonic looping from northward
to eastward to southward to eastward again as it exits the Gulf of Mexico through the Straits of Florida. This clockwise turning of the Gulf Loop Current is part of a cycle of growth (penetrating into the Gulf of Mexico almost to the latitude of the Mississippi River Delta, fE301N), then turning eastward and southward to run along the west Florida escarpment. Near the latitude of Key West (fE241N), the current turns sharply cyclonically and begins to run eastward in the Straits of Florida, where it is now called the Florida Current. Once the Gulf Loop Current has reached its maximum latitudinal extent, a large anticyclonic current ring, 100– 150 km radius, separates from the flow, and the main current reforms farther south near the latitude of the Florida Keys. In its southernmost position, the Gulf Loop Current flows into the Gulf of Mexico, and turns rather sharply in an anticyclonic turn to exit almost directly into the Straits of Florida (Figure 4). Anticyclonic Gulf Loop Current rings have all the T–S and flow features of the Gulf Stream system, just as in the Yucatan Channel and in the Straits of Florida. The ageostrophic dynamic balance is 7n 2 =r þ f n 1=r @p=@ n ¼ 0, where r is the radius of curvature, and where 7n 2 =r is positive in cyclonic
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Figure 4 Sea surface height (h) in meters from numerical model calculations as reported in Mooers and Maul (1998). Geostrophic surface currents are calculated from f n ¼ 1=r@p=@ n ¼ g@h=@ n . Anticyclonic eddies are isolated concentric height maxima, the largest of which is shown in the western Gulf of Mexico; cyclonic eddies have the opposite surface height field.
curvature and negative in anticyclonic. Accordingly, flow is super-geostrophic in anticyclonic turns, and sub-geostrophic in cyclonic turns. Identical dynamics describe midlatitude upper tropospheric flows in Earth’s atmosphere, notably in the Jet Stream. A separated Gulf Loop Current ring travels westward into the western Gulf of Mexico, most probably by a self-propulsion mechanism associated with the beta effect, b ¼ @f =@y, of differing Coriolis parameter between the southern and northern ring edges. Using the hydrostatic equation, @p ¼ rg@z, the horizontal pressure gradient term 1=r @p=@ n may be written as g@h=@ n, and for an anticyclonic Gulf Loop Current ring with a diameter of 300 km and @h ¼ 0:75 m, it is calculated using n 2 =r þ f n g@h=@ n ¼ 0 that eddies self-propagate at speeds of 5–10 cm s1 ðE5 10 km d1). Direct observations from satellite-tracked buoys and from satellite altimeter measurements of sea surface height (h) substantially agree with such calculations. As these Gulf Loop Current rings travel to the west, they begin to spin down, losing their momentum per unit volume, rn, to horizontal friction and eventually mixing in with the ambient Gulf of Mexico Common Water. The lifetime of the current rings is typically 6 months, and they carry with them
the temperature, salinity, and other characteristics of their source region, the Caribbean Sea. Approximately 3 1013 kg y1 of salt is injected into the western Gulf of Mexico by an average Gulf Loop Current ring. Thus the salt balance of the western Gulf of Mexico is decidedly nonMediterranean, E þ d ¼ P þ R, because in the IAS the classical evaporation/precipitation/runoff equation requires an additional term d to account for the infusion of high salinity water from the rings. Typical values for these terms are E2P E35 cm y1, R E75 cm y1, and the volume of fresh water to maintain the salt balance d E40 cm y1. The Gulf Loop Current interacts with the fourth great riverine system of the IAS, the Mississippi River. As the Gulf Loop Current nears the Mississippi Delta, it is observed to entrain or advect the river water to the east. Mississippi River water has been observed by its low surface salinity all along the eastern edge of the Current, into the Straits of Florida, and up the east coast of the USA at least to Georgia (fE321N), the northern boundary of the IAS. Thus the four great rivers of the IAS, the Amazon, the Orinoco, the Magdalena, and the Mississippi, and the many smaller tributaries, are all known to interact with the oceanic flows and to be
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INTRA-AMERICAS SEA
carried great distances by them. This is an important transport mechanism in the IAS whereby riverine flotsam and jetsam is found on distant shores. This same flow-through regime is also responsible for the considerable impact of tars from maritime commerce and from oil drilling on the highly valuable tourist beaches of the area (cf . Figure 3). Gulf Loop Current ring shedding seems to have a cycle of 10–11 months on average, with some rings being shed in as few as 6 months and others taking 17 months. This is a surprising frequency since it is not at the annual harmonic where many other oceanic features have a spectral peak. Gulf Loop Current ring shedding has been simulated by IAS numerical models, and super-annual periods are often calculated. The cycle of ring formation does not seem to be forced by the R Runmistakable annual cycle of volume transport, nðx; zÞdx dz in the Gulf Stream system, with its maximum in June and minimum in October, nor by variability in the Caribbean Current along 151N, which has a spectral peak at about 75 days. Connectivity between flow variability in the Caribbean Sea and the Gulf of Mexico seems remarkably weak. In the Straits of Florida, the Florida Current turns cyclonically as it passes between Cuba, Florida, and the Bahamas. The lesser passages of the Straits of Florida contribute small amounts to the total volume transport, which by now is at the level of 30 G s1 (30 106 m3 s1) or about 30 Sv. At the latitude of Palm Beach, Florida (fE271N), the meridional RR oceanic heat transport, rCpTnðx; zÞdx dz, is about 1.5 petawatts (1.5 1015W), an amount comparable with the atmosphere at the same latitude. Transport is known to vary on timescales ranging from days to years. Fortnightly volume transport changes are observed to range from 20 Sv to 40 Sv, a value much larger than that of the annual cycle, which is more likely less than 75 Sv. In the narrow confines of the northern Straits of Florida, lateral friction is significantly larger than in the open sea. Here, from extensive in situ studies, the Guldberg-Mohn friction coefficient J in quasi-geostrophic flow fv ¼ g@h=@x þ Jv has been estimated at approximately 4 105 s1, one to two orders of magnitude larger than away from continental boundaries. In addition, the Florida Current axis meanders to the west when transport is high and to the east when it is less. Such detailed transport variations in other regions of the IAS are not as well documented as in the Straits of Florida between Palm Beach and the Bahamas. North of the Straits of Florida the current is called the Gulf Stream, a name it retains to the offing of Nova Scotia. The flow northward to Cape Hatteras,
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North Carolina (fE351N), seems to be topographically controlled, whereby H in the potential vorticity equation (above) directs onshore and offshore meandering. Near Charleston, South Carolina (fE321N), a notably shallow area decreases H significantly, and the current responds by turning anticyclonically ðBo0Þ into deeper water. Once offshore, the larger value of H causes the flow to turn cyclonically again ðB > 0Þ in a series of meanders as it progresses downstream (cf . Figure 3). Along the outer boundary of the IAS, a surface flow with a great deal of changeability is observed, called the Antilles Current. This intermittent current, carrying on average about 15 Sv, progresses northward up the margin of the Caribbean Sea, along the eastern outer banks of the Bahamas, and eventually joins the Gulf Stream north of Cape Canaveral, Florida (fE291N). Satellite-tracked buoys and subsurface floats both show the latitude of the Bahamas to be an area of great temporal and spatial variability, with many eddies of various sizes, and with inflows to the Straits of Florida through the Old Bahama Channel and the Northwest Providence Channel. The general sense is that of converging surface flows all feeding the Gulf Stream system.
Subsurface Flow Regime The strong surface currents of the Gulf Stream system decrease with depth. Mathematically, this can be explored by applying Leibnitz’ rule to R z the integral form of the hydrostatic equation p ¼ h rgdz . The geostrophic equation (above) can then be expressed as: Z z g@r=@ ndz rf v ¼ rg@h=@ n þ h
where the first term on the right-hand side is the barotropic term, and the integral term on the righthand side is the baroclinic term. Facing downstream @r=@ n is negative, and the surface current (at z ¼ h) decreases with depth until the two terms on the righthand side become equal and opposite. This depth is called the level-of-no-motion and n ¼ 0. In the Yucatan Channel, the level-of-no-motion is approximately 1200 m, but in the northern Straits of Florida (fE271N) northward-flowing currents as much as 0.3 m s1 reach to the seafloor at 800 m. Details of the flows into and out of the IntraAmericas Sea at depth in other passages are less well known than the surface flows. Numerical models and observations suggest a general inflow though the passages of the Lesser and Greater Antilles into the Caribbean Sea, an outflow through the Yucatan Channel into the Gulf of Mexico, and continuing flow into the North Atlantic Ocean north of the
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CURRENT EDGE
VORTEX STREET
YUCATAN PENINSULA
CLOUD COZUMEL ISLAND
Figure 5 LANDSAT negative image of surface-wave glitter patterns showing von Karman vortices downstream of Cozumel Island in the Yucatan Channel. For scaling, Cozumel Island is approximately 50 km long.
Bahamas. Near the sill of the Yucatan Channel, approximately 100 m above the ocean floor, the flow is decidedly from the Gulf of Mexico into the Caribbean Sea. In the Windward Passage, there is also evidence of north-eastward outflow at depth, but it seems not to be as persistent as that in the Yucatan Channel. A major characteristic of the deep waters of the IAS is their nearly isothermal and isohaline profiles below the depth of the major sills. The ocean, being a stratified fluid, tends to inhibit vertical mixing. Thus the sub-sill depth waters are characterized by nearzero vertical density gradients, @r=@zE0, and are neutrally stable. Deep IAS waters have the T–S characteristics of the offshore waters of the juxtaposed North Atlantic Ocean, which seem to spill over the sills from time to time to replenish and ventilate the water interior to the IAS. Thus there must be a surging of sorts to bring into the Caribbean basin in particular, renewing mid-depth North Atlantic Common Water. Along the eastern margin of the IAS at about 2000–3000 m water depth is a southward-flowing mid-level current called the Deep Western Boundary
Current (DWBC). The DWBC has its genesis in the thermohaline circulation of the North Atlantic Ocean, north of the Denmark Strait, and is part of the global conveyor belt. The DWBC flows along the entire outer boundary of the IAS, and has a volume transport of approximately 15 Sv. This extremely important flow is climatically linked to the role of the ocean in Earth’s heat budget and may participate in the complexities of the IAS’s deeper flow patterns.
Other Currents in the Intra-Americas Sea Tidal currents in the IAS are generally weaker than in other semi-enclosed seas. The tides are typically semi-diurnal on the Atlantic Ocean margin of the IAS (M2 and S2 constituents usually), and progressively become diurnal in the Gulf of Mexico where the K1 and O1 constituents dominate. Estuaries such as the Mississippi Delta are of the saltwedge category, mostly because the tidal currents and ranges are small and the river flows very large
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(average for the Mississippi River is about 103 km3 y1E0.03 Sv). Tidal currents around many IAS islands are similarly weak, with extremes rarely exceeding n ¼ 1 m s1 even in passes through the many bar-built barrier island lagoons. Inertial currents are a ubiquitous feature of the ocean, and are characterized by periods n ¼ 12h =sin f. In the northern IAS, inertial currents often have periods equal to the dominant diurnal tidal currents, such as the K1 or the O1 because sinfE0.5. At these critical latitudes, the inertial currents are not separable from the diurnal currents in the tidal spectrum. Inertial currents, when they occur, are intermittent and have velocities typically below n ¼ 0.2 m s1. Current flows past islands can induce complex patterns in their lee. Numerical models and observations suggest von Karman vortex streets downstream of many island land masses (Figure 5). Such complex currents can cause engineering design complexities, particularly regarding waste disposal and spills. Similarly, with the normally low wave heights so characteristic of the IAS, the longshore and littoral currents are also weak, although in
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certain areas, especially the east coast of Florida, dangerous wave-induced rip currents are very common. Perhaps the most significant physical marine hazard in the Intra-Americas Sea is the combined storm surge and inverted barometer effect associated with hurricanes. Along linear coasts, the water level elevation from a major storm can exceed 7 m, on top of which may be 3 m or greater wind waves. Little is directly known of the currents associated with storm surges, but indirect evidence suggests that they can exceed n ¼ 1–2 m s1, especially in the vicinity of harbor entrances and inlets. Small islands are less at risk from large storm surge-driven currents than are long, low coasts with shallow offshore bottom topography. While storm surge currents are transient features of the IAS with timescales of approximately half a day, they can be costly to infrastructure, and very dangerous to human life. Upwelling is a vertical current of note in the IAS, especially along the long east–west tending north coast of South America. Here the zonal wind stress with mass transtx can force an Ekman upwelling R port per unit width My ¼ rvdz given by My ¼ tx =f.
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t N E
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Figure 6 Summary of surface currents in the Intra-Americas Sea. Maximum IAS sea surface height variability h ¼7 24 cm is centered in the Gulf Loop Current (GLC) at f ¼ 261N, l ¼ 881W; a second maximum in the Caribbean Current of h ¼ 7 12 cm is centered at f ¼ 151N, l ¼ 771W.
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Accordingly, the easterly Trade Winds force a north- g gravity ward mass transport along the coast, and the ocean h sea surface height responds with classical coastal upwelling as seen in n direction vector parallel to pressure the 31C cooler sea surface temperatures along Venegradient zuela. The same physical circumstances cause open- p pressure ocean Ekman pumping in the wake of hurricanes. The r radius of curvature hurricane’s large cyclonic wind stress, t ¼ rair Cd v2air , t time forces mass transport Mx;y in all directions away from u,v,w eastward, northward, upward speed the storm center with attendant lifting of the ther- n velocity vector orthogonal to n mocline and upwelling. Lower sea surface tempera- x, y, z east, north, vertical Cartesian tures are often observed as a cool streak in the wake coordinates air-sea drag coefRcient of these intense air–sea storm systems. Cd specific heat While the danger from seismic sea waves (tsu- Cp evaporation nami) is recognized as another although largely E gigaliters unappreciated natural hazard of the IAS, it is the Gl water depth waves and wave particle motions that create cur- H mass transport rents of such great danger. Caribbean tsunami M precipitation waves have been observed to exceed 9 m in height. P R river runoff Since a tsunami is a progressive shallow-water pffiffiffiffiffiffiffi salinity wave with celerity c ¼ gH, the maximum S Sverdrups currents come at the wave crest and at the wave Sn temperature trough. These currents probably exceed v ¼ T 10 m s1 , and have timescales of several minutes. In W watts that short amount of time however, even more d salinity anomaly from eddies danger exists than with storm surge currents, and f latitude the small islands are equally as vulnerable as are l longitude the continental coasts. r density z vorticity t wind stress Conclusions O Earth’s rate of angular rotation Intra-Americas Sea surface currents (Figure 6) are dominated by a single fact: the IAS is the formation region of the Gulf stream system. Except along the northern coast of the Gulf of Mexico, the volume See also transport of the interior thermohaline component Sphenisciformes. Tides. of IAS currents is minuscule compared with the wind-driven component. While there are important external and peripheral currents associated with the global thermohaline flow such as the Deep Western Further Reading Boundary Current, it is the North Brazil Current– Guianas Current–Caribbean Current–Yucatan Cur- Gallegos A (1996) Descriptive physical oceanography of the Caribbean Sea. In: Maul GA (ed.) Small Islands: rent–Gulf Loop Current–Florida Current–Gulf Marine Science and Sustainable Development. Stream family of atmospheric wind stress-forced Washington: American Geophysical Union. advective movements that characterizes the region Maul GA (ed.) (1993) Climatic Change in the Intra(cf . Figure 4). All these ‘currents’ are in reality one Americas Sea. London: Edward Arnold. current that, coupled with air–sea heat and mois- Mooers CNK and Maul GA (1998) Intra-Americas Sea ture fluxes and winds, integrate into a single conCirculation. In: Robinson AR and Brink KH (eds.) The tinuum that connect the Intra-Americas Sea and its Sea, Vol. 11. New York: John Wiley & Son. peoples. Murphy SJ, Hurlburt HH, and O’Brien JJ (1999) The
List of Symbols c f
wave celerity Coriolis parameter
connectivity of eddy variability in the Caribbean Sea, the Gulf of Mexico, and the Atlantic Ocean. Journal of Geophysical Research 104: 1431--1453. Schmitz WJ Jr (1995) On the interbasin-scale thermohaline circulation. Reviews in Geophysics 33: 151--173.
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INTRUSIONS Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1363–1367, & 2001, Elsevier Ltd.
Introduction In most frontal regions, where waters of different salinities and temperatures meet laterally, an interleaving of the different waters is observed. These features are commonly referred to as intrusions. Sometimes a single layer of water from one region is advected into the other region, such as the Mediterranean salt tongue in the North Atlantic, by either a mean flow or eddy motion. Multiple layers of the two different water masses are also seen quite often. The driving mechanism for these multiple intrusions is related to horizontal gradients in salinity and temperature and the small-scale (e.g., smaller than the thickness of the intrusions) mixing occurring between the interleaving layers. In this article, only intrusions produced by this latter process are discussed. Both observational and theoretical studies are presented. Frontal regions are locations where waters of different temperature and salinity meet and interact. They are usually characterized by relatively large horizontal gradients in these two properties. Fronts have been found in the coastal ocean, at the shelfbreak and at the boundaries of major currents, such as the Gulf Stream and Antarctic Circumpolar Current. An example of a front (Figure 1) is shown by the azimuthally averaged salinity structure of a Mediterranean eddy. A Mediterranean eddy (Meddy) is a coherent eddy of Mediterranean Sea water found in the eastern North Atlantic Ocean. The front with its larger horizontal gradients in salinity is located at a depth range of 700–1300 m and with a radius of 15–30 km. The temperature field has a similar structure to the salinity structure shown in Figure 1. With the horizontal change in salinity, it would be expected that there would be a horizontal change in the density of the sea water. However, the effect of the horizontal change of temperature on the density nearly completely compensates the density change due to the salinity change across the front. Thus, the density surfaces are nearly horizontal. However, there is a slight upward (downward) tilt of density surfaces in the lower (upper) half of the Meddy. The resulting pressure gradient balances the geostrophic
flow of the eddy. Along-front geostrophic flows are found at most fronts. A closer look at the structure of temperature and salinity in the frontal region shows an interleaving of water with the characteristics of the temperature and salinity on the two sides of the front. The temperature and salinity of the water in these interleaving layers show evidence that mixing of the two water types has also occurred. Figure 2 shows a section of closely spaced (e.g., 1–2 km) vertical profiles of salinity starting from the center of the Meddy (Figure 1) and moving towards the outside edge. In each profile, there are wiggles in the salinity field, typically of 1–2 km vertical scale, which represents the water moving horizontally from the center of the Meddy to the edge or vice versa. These wiggles are referred to as intrusions. Intrusions like these are found in most frontal regions such as those associated with the Gulf Stream and Antarctic Circumpolar Current. The observed interleaving of temperature and salinity is thought to develop as an instability of the thermohaline front. These fluctuations lead to regions of enhanced double-diffusive mixing. Two types of double-diffusive mixing can occur: salt-fingering under the warm, salty layers and diffusiveconvection under layers that are relatively cold and
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Figure 1 Azimuthally averaged cross-section of the salinity of a Mediterranean eddy (Meddy) embedded in eastern North Atlantic water. This survey, the second one of this Meddy, was made in June 1985, PSU, practical salinity units.
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Density flux
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Cold and fresh
Salt-fingering interface Diffusive-convection interface
Frontal region Figure 3 A schematic of the interleaving layers representing the intrusions. The open arrows indicate the cross-frontal motion driven by the depth-varying density flux (solid arrows). In this diagram, salt-fingering is the dominant form of double-diffusion; thus, the warm, salty water rises as it crosses the front.
fresh (Figure 3). Both forms of double-diffusive mixing generate a downward density flux, that is, a release of potential energy. The convergence or divergence of this density flux makes the intrusion either heavier or lighter, respectively. These density changes produce pressure gradients which drive the interleaving motions across the front. If the density
flux of salt-fingering exceeds that of diffusive-convection, waters in the warm, salty layers become less dense and, therefore, rise as they cross the front. The cold, salty layers become more dense and sink as they cross the front. It is believed that this case applies for the intrusions found for the lower half of the Meddy. If diffusive-convection dominates (which is believed
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INTRUSIONS
to be the case for the intrusions occurring in the upper half of the Meddy), water in the cold, fresh layers should rise across the front and the warm, salty layers sink.
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There have been many observations of intrusions in vertical profiles of salinity and temperature taken in frontal regions. Other than demonstrating the presence of intrusions and indicating their vertical scale, it is difficult to make any other conclusions about the dynamics of the intrusions. Closely spaced profiles (e.g., Figure 2) show that the horizontal structure of intrusions is complex. Although it is possible to track an intrusion across several kilometers (and several profiles), the structure of the individual intrusion changes significantly. In addition, some intrusions appear to start and end abruptly. One of the problems of interpreting this type of data is that the frontal region usually has a horizontal velocity field associated with it. Although the water in the intrusions is moving across the front, it is also being advected along the front by the geostrophic current of the front. Therefore, some of the observed crossfrontal variability could be due to differential advection of the intrusions along the front. Most of the observations of intrusions have been single surveys of the front; the evolution and the dynamics of the intrusions could not be determined. Even if multiple surveys are undertaken, temporal changes in the intrusions cannot be separated from possible alongfrontal variations of the intrusions. However, there has been one study where some of the dynamics of the intrusions could be investigated. This was an experiment to determine the evolution of a Mediterranean eddy. The front (Figure 1) between these two water masses can be thought of as a circular front. Thus, the problem of differential along-front advection of the intrusions is removed since the front loops back on itself. It would be expected that the individual cross-frontal transects could be typical for all radial sections and that alongfrontal variations are small. Thus, the cross-frontal transects could be used to determine the intrusion dynamics. This Meddy was surveyed four times over a two-year period as it decayed. The vertical structure of the intrusions evolves as the cross-frontal temperature and salinity gradients change (Figure 4). For the first year of the study, the Meddy had a core region unaffected by intrusive mixing. During this time, the intrusions appeared to have a similar wiggly vertical structure at all locations for both surveys with vertical scale of about
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Figure 4 Vertical profiles of salinity made through the intrusive region of the Meddy during four surveys: October 1984 (solid line), June 1985 (dashed line), October 1985 (dotted line) and October 1986 (bold line). The profiles have been offset by 0.2 PSU (practical salinity units) from each other.
20 m. The wiggles are rather smooth (i.e., sinusoidal) and have approximately the same vertical scale. By the time of the third survey, the intrusions just reached the center of the Meddy. We can imagine that there was a constant cross-frontal gradient (driving the intrusions) for the first year of observation and that the intrusions had passed their initial (exponential) growth stage. By the time of the third survey, some of the intrusions appeared to have a step-like structure. A year later, the intrusions had a more pronounced step-like structure with a larger vertical scale, about 50 m (Figure 4). This structure is probably representative of decaying intrusions. Double-diffusive processes were still active vertically but the horizontal advection mean gradients were less important; there was not a supply of new water unaffected by the mixing. One major question that remains concerns the cross-frontal fluxes of heat and salt by the intrusions. In order to address this question, it is necessary to measure the very weak velocities in the intrusive layers or observe the large-scale changes in the
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properties of the frontal region. For the first year of study of the Meddy, it had a core region with very little horizontal variability (Figure 1). Using the rate at which the intrusions moved into this central core region, extremely small cross-frontal velocities, u0 , on the order of 1 mm s1 were found. Using the salinity anomalies associated with the intrusions, S0 , and this order of magnitude estimate for the crossfrontal velocity, the average cross-frontal flux of salt, FS ¼ /u0 S0 S, was calculated. Parameterizing this flux in terms of horizontal diffusion, FS ¼ K H
1 qS r qr
½1
where KH is the horizontal eddy diffusivity and qS/qr is the mean horizontal (radial) salinity gradient across the front; an eddy diffusivity coefficient of 0.4 m2 s1 was found. The dominant mechanism responsible for the decay and eventual demise of the Meddy was thermohaline interleaving, presumably driven by double-diffusive buoyancy fluxes. Over the 2-year observation period, intrusions at the edge of the Meddy core eroded the warm and salty central region from an initial diameter of 60 km until the core was no longer detectable. Using the rate at which the salinity and temperature of the Meddy at a specific radius changed, an eddy diffusivity could be estimated. qS q2 S 1 qS þ ¼ KH qt qr2 r qr
! ½2
Likewise, by integrating eqn [2] from the center to a specified radius, changes in the salt and heat content of the Meddy can be used to estimate an eddy diffusivity. It was found that an eddy diffusivity of 1–5 m2 s1 could be used to parameterize the crossfrontal fluxes of the intrusions. To date, this Meddy study has been the only one where estimates of the fluxes by intrusions could be made. Estimates of the horizontal eddy diffusivity ranged from 0.5 to 5 m2 s1. An attempt was made to understand the dynamics of the intrusions with these surveys but the temporal sampling (six months) was too infrequent to be of use in investigating the evolution of the intrusions.
Theoretical Studies The driving mechanism for the cross-frontal velocity of intrusions (i.e., the horizontal pressure gradients) is due to divergences in the vertical density fluxes, generally assumed to be due to double-diffusive mixing (Figure 3). Most of the theoretical studies to
date have looked at the initial growth of the intrusions using linear stability analysis with parameterizations for the vertical flux by salt-fingers. In these linear stability calculations, the background frontal structure is assumed to have linear gradients, both horizontally and vertically, of salinity and temperature. The horizontal gradients are chosen such that there is no horizontal gradient in density and thus, no along-front velocity. Vertical gradients are chosen such that the background structure is unstable to double-diffusion, usually saltfingering. Double-diffusive mixing is parameterized as a constant eddy diffusivity for salt (or heat) and a constant ratio of the heat to salt flux. The perturbations are assumed to be small, so there are no inversions in temperature and salinity. The linear stability analysis predicts the vertical scale, crossfrontal and along-frontal slopes of the fastest-growing unstable mode given the salinity and temperature gradients. These properties have been compared to observations and have shown general agreement. For typical horizontal gradients of salinity and temperature found in frontal regions, the growth rate of the fastest mode has an e-folding timescale on the order of 10 days. Inclusion of a background velocity shear due to the sloping isopycnals across the front can produce faster-growing intrusions with an e-folding timescale on the order of several days. Linear stability studies predict properties of the initial growth stage, in which fluxes grow exponentially, but say nothing about the finite amplitude ‘steady’ state properties. When the intrusions reach finite amplitude, the fluxes of heat and salt by the interleaving should reach a constant value. Since fronts in the ocean exist much longer than the time for the intrusions to grow, intrusions spend most of their lives in the finite-amplitude state. Therefore, the usefulness of extrapolating intrusion properties and fluxes from linear theory is questionable. For growing intrusions to reach an equilibrium, a three-way balance between salt-finger, diffusiveconvection and (cross-frontal) advective fluxes is necessary. The initial instability may set the vertical scale of the finite-amplitude intrusions, but the crossfrontal fluxes may depend critically on the form of the equilibrium that the growing intrusions eventually reach. A numerical model verified that small amplitude intrusions, predicted by linear stability analysis, evolved into large amplitude, equilibrium, intrusions. When the amplitude of the intrusion becomes large enough that temperature and salinity inversions occur, the growth of the intrusion slows and reaches an equilibrium state. This equilibrium state is characterized by interleaving layers with saltfingering and diffusive-convection occurring at the
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interfaces separating statically unstable ‘convecting’ layers. As expected, the three-way flux balance is achieved. As well as obtaining a balance in the advection and mixing of the salinity and temperature, there must be a momentum (energy) balance. The double-diffusion mixing lowers the potential energy of the system. This potential energy is converted into kinetic energy within the convecting layers. In addition, the convecting layers allow a large flux of momentum from the salt-fingering interface to the double-diffusive interface. The friction between the interleaving layers balances the pressure gradient produced by the density flux divergence.
dimensional structure of intrusions to evaluate the two-dimensional studies done to date.
Summary
Further Reading
The presence of intrusions in frontal regions has led oceanographers to believe that they must be important in the cross-frontal fluxes of heat and salt. However, at present, these fluxes are almost impossible to observe in the ocean. Thus, we must rely on theoretical and numerical studies to address this important question. In order to be useful, these studies must predict properties of intrusions which can be compared to observations. To date, comparisons have been limited to the vertical length scale of intrusions from single vertical profiles of temperature and salinity. Predictions of the slope of the intrusions (relative to density surfaces) in the crossfrontal and along-frontal directions have been compared to the few cross-frontal sections made. With improvements in navigation with global positioning satellites and the advent of undulating towed bodies, rapid three-dimensional high-resolution mapping of intrusions can be undertaken. Future work, both numerical and observational, will use the three-
See also Double-Diffusive Convection. Meddies and SubSurface Eddies. Shelf Sea and Shelf Slope Fronts. Upper Ocean Mean Horizontal Structure. Upper Ocean Mixing Processes. Water Types and Water Masses.
Hebert D, Oakey N, and Ruddick B (1990) Evolution of a Mediterranean salt lens: scalar properties. Journal of Physical Oceanography 20: 1468--1483. May BD and Kelley DE (1997) Effect of baroclinicity on double-diffusive interleaving. Journal of Physical Oceanography 27: 1997--2008. McDougall TJ (1985) Double-diffusive interleaving. Part II: Finite amplitude, steady state interleaving. Journal of Physical Oceanography 15: 1542--1556. Ruddick B (1992) Intrusive mixing in a Mediterranean salt lens – intrusion slopes and dynamical mechanisms. Journal of Physical Oceanography 22: 1274--1285. Ruddick BR and Hebert D (1988) The mixing of Meddy ‘Sharon’. In: Nihoul JCJ and Jamart BM (eds.) Small-Scale Mixing in the Ocean. Elsevier Oceanography Series, vol. 46. Amsterdam: Elsevier. Toole JM and Georgi DT (1981) On the dynamics and effects of double-diffusively driven intrusions. Progress in Oceanography 10: 123--145. Walsh D and Ruddick B (1998) Nonlinear equilibration of thermohaline intrusions, Journal of Physical Oceanography 28: 1043--1070.
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INVERSE MODELING OF TRACERS AND NUTRIENTS R. Schlitzer, Alfred Wegener Institute for Polar and Marine Research, Bremerhaven, Germany & 2009 Elsevier Ltd. All rights reserved.
Introduction Seawater contains a large variety of dissolved and particulate constituents, commonly referred to as tracers. Many of these tracers (such as the dissolved nutrients, phosphate, nitrate, and silicate, or the radioactive carbon isotope 14C (radiocarbon)) are natural and are part of the ocean system since geological times. Other tracers, such as the chlorofluorocarbons (CFCs) and various radioactive isotopes from nuclear bomb testing, are of anthropogenic origin, and have been introduced into the ocean only during the last decades. The distributions of anthropogenic tracers in the ocean change markedly on decadal or annual timescales. Naturally occurring tracers, like the nutrients, are believed to be near steady state. However, recent data indicate decadal and inter-annual changes in ocean oxygen and nutrient distributions, likely caused by changes in circulation and biogeochemical processes associated with global climate change.
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Some ocean tracers, such as radiocarbon and the CFCs, act as passive tracers. They are chemically inert and are being transported and redistributed in the ocean by currents after entering the ocean from the atmosphere. When dense surface waters sink and spread in the ocean interior, they carry the tracer signal with them and produce tongues of high tracer concentrations along their spreading paths. An example of such tracer tongues and concentration maxima in the ocean interior can be seen in the zonal CFC section in Figure 1. Note the pronounced CFC maxima at about 1500- and 3500-m depth at the western boundary, revealing the two cores of the deep western boundary current (DWBC) carrying tracer laden waters southward at the western side of the basin. The spatial extent of the tongues and the concentration gradients along the tongues can, in principle, be used to determine the pathways and velocities of the currents. Such inferences require advanced mathematical methods, some of which are described below. Other tracers (like the macronutrients phosphate, nitrate, and silicate, or the micronutrients Fe and Zn) are chemically and biologically active and are a prerequisite for biological production in the ocean’s surface layer. Surface nutrient concentrations are very low in most regions of the world ocean due to
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Figure 1 Zonal section of chlorofluorocarbon 11 at about 241 N in the Atlantic, clearly showing the traces of the upper and lower branches of the deep western boundary current (DWBC). Chlorofluorocarbons (CFCs) are man-made gases that invade the ocean from the atmosphere in increasing amounts since the 1940s. Before that time the ocean was CFC-free. High CFC concentrations in the ocean interior are clear indications of vigorous subsurface currents.
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INVERSE MODELING OF TRACERS AND NUTRIENTS
biological utilization. Biologically produced organic particles as well as calcium carbonate and opal shells sink and dissolve in deeper parts of the water column, thereby returning nutrients to the dissolved pool, while utilizing and drawing down dissolved oxygen. Overall, the biogeochemical processes act as a vertical nutrient and carbon pump creating significant vertical gradients and pronounced subsurface
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nutrient maxima at the depths where particle remineralization occurs. Subsurface nutrient maxima are often correlated with oxygen minima, clearly indicating the remineralization of organic material. The signatures of these processes in the nutrient and oxygen distributions are large and easily detected (Figure 2) and reveal information about the underlying processes. The strength of the biological
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Figure 2 Meridional sections of (a) salinity, (b) phosphate, and (c) oxygen along WOCE section A16 in the Atlantic. In addition to ocean circulation, the phosphate and oxygen distributions are also affected by biological nutrient utilization near the surface and by subsurface remineralization of sinking organic material. Organic matter remineralization releases dissolved nutrients while utilizing dissolved oxygen, and can be clearly identified between 200- and 1500-m depth in the tropical and subtropical areas.
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production, for instance, can be inferred from the observed strength of the vertical nutrient gradient, and the depth range of particle remineralization can be determined from the observed vertical position and spatial extent of the nutrient maxima and oxygen minima. Differences in the vertical structure of different nutrients (e.g., phosphate, silicate, carbon, and alkalinity) reveal differences in the remineralization depths of organic material, CaCO3, and opal. Basin-wide observations of a variety of oceanic tracers have been conducted since the 1950s. The first coordinated and global tracer program, GEOSECS, produced tracer data of unprecedented quality and coverage during the 1970s. More recently, the World Ocean Circulation Experiment (WOCE) has provided an even more detailed tracer data set describing the distributions during the 1990s. These data are publicly available in electronic form over the Internet or as colored distribution plots in printed atlases. Availability of original tracer data is essential for the inverse methods described below.
Inverse Model Concepts Deriving quantitative results about the underlying biogeochemical processes from oceanic nutrient and tracer data, and separating the effects of biogeochemistry from the effects of circulation is a challenging task, and requires the use of coupled physical/biogeochemical numerical models. There are a variety of possible approaches, which can broadly be divided into two categories. The so-called ‘forward models’ assume rates of physical and biogeochemical processes to be known a priori, and require ocean currents (or the physical forcing at the ocean surface), as well as biological production and particle remineralization rates to be specified as input. Oceanic tracer concentrations are treated as unknowns, and the tracer fields that would evolve under the assumed flows and biogeochemical parameters are then simulated. Forward models, in general, lead to mathematical systems that are relatively easy to solve. The conceptual disadvantage of forward models is that physical and biogeochemical rate information that is supposed to be determined from the tracer data and, in fact, only becomes available after the data evaluation, is required a priori to enable the model run. ‘Inverse models’, in contrast, follow a seemingly more intuitive approach, treating the measured tracer concentrations and other auxiliary knowledge formally as knowns, whereas the physical and biogeochemical parameters, to be determined on the basis of the tracer data, are treated as unknowns.
As described in more detail below, the inverse approach generally leads to underdetermined mathematical systems that are much harder to solve than the systems encountered in forward models. Error and resolution analysis are two issues of particular importance when solving underdetermined inverse problems. First, the solved-for physical and biogeochemical parameters depend directly on the tracer data, and errors in the data propagate into errors in the solution. Second, owing to the incompleteness of information in underdetermined systems, the unknowns are usually not fully resolved. Instead, only specific linear combinations of unknowns may be well constrained by the data, while individual unknowns or other combinations of unknowns may remain poorly determined. Both, error and resolution analysis are essential for a quality assessment of the solution of underdetermined systems. Two widely used practical inverse approaches are presented in the next two sections (‘Estimating absolute velocities and nutrient fluxes across sections’ and ‘Estimating carbon export fluxes with the adjoint method’). These serve as examples to describe details of the mathematical methods and to list achievable results. The first method, the section inverse approach, infers nutrient, carbon, and tracer fluxes across sets of sections, based on hydrographic, tracer, and nutrient data along these sections. The second example describes an application of the adjoint method for the determination of ocean currents, biological productivity, and downward particle fluxes. This method is specifically adapted for the utilization of many different tracers and can handle problems with heterogeneous and sparse data coverage.
Estimating Absolute Velocities and Nutrient Fluxes across Sections The section inverse approach developed by Wunsch exploits hydrographic data along sets of oceanographic sections, and allows estimation of absolute flow velocities perpendicular to the sections. The method was later extended to include nutrient and oxygen data allowing to estimate nutrient and oxygen fluxes across the sections. The section inverse method is based on the geostrophic principle, which for a given pair of hydrographic stations allows calculating the geostrophic velocity vg(z) perpendicular to the connecting line between the two stations as function of depth z. A reference velocity b has to be added to the geostrophic velocity to obtain the absolute flow velocity
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vðzÞ ¼ vg ðzÞ þ b
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INVERSE MODELING OF TRACERS AND NUTRIENTS
at a given station pair. The reference velocity b is left unknown by the geostrophic calculation. Determining these unknowns was a major challenge in physical oceanography for many decades. There is one unknown reference velocity bi for every station pair in the sections considered. Thus, for global networks consisting of hundreds of sections, as in the case of the WOCE survey, the number of unknown velocities bi may amount to several thousands. Steady-state mass and tracer conservation equations for all subdomains formed by the intersecting hydrographic sections (see Figure 3) are written as constraint equations for the unknown bi. For a given domain the conservation equation of tracer C is expressed as X ðc¯ v¯ gi þ c¯ bi Þ di Dxi ¼ 0
½2
i
where the summation is over all station pairs along the boundary of the domain, c¯ v¯ gi is the vertically averaged tracer flux density arising from the
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geostrophic flow, c¯ is the vertically averaged tracer concentration, di is the mean water depth of pair i, and Dxi is the distance between the two stations of the pair. Mixing terms are ignored for the sake of simplicity. Mass budget equations are obtained by replacing c with density r. It is important to note that the budget equation [2] is linear in the unknown bi, and that the known components of the fluxes involving the geostrophic velocities can be moved to the right-hand side. Formulation of conservation equations for all domains (whole water column and suitably defined layers) and for all tracers considered results in a set of linear equations Ab ¼ G
½3
where b is the vector of m unknown reference velocities (m is the total number of station pairs), A is an n m matrix containing the coefficients of the n budget equations, and P G is an n vector with the known right-hand sides i c¯ v¯ gi di Dxi of the budget
(a)
T1 d1
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Figure 3 Simple examples of (a) a three-station pair defining sections I and II, and (b) the station configuration and bottom depths di in sections I and II, where dashed lines are supposed layer interfaces. Reproduced from Wunsch C (1978) The North Atlantic general circulation west of 501 W determined by inverse methods. Reviews of Geophysics 16: 583–620, with permission from the American Geophysical Union.
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equations. The objective is to solve the system of linear equations for b. Once this is achieved, the absolute flow velocities can be calculated according to [1], and the fluxes of any tracer across a given cross section can be obtained by integrating c v over the interface. Finding the solution b of [3] is complicated for two reasons: (1) the number of equations n is typically much smaller than the number of unknowns m, leaving [3] underdetermined, and (2) the right-hand side G is based on hydrographic measurements and therefore contains errors. Because of the errors in G, there is no exact solution to [3], in general, and one has to resort to the least-squares solution that minimizes the residuals r ¼ Ab þ G. It should be noted that row and column weighting of [3] is usually applied before calculating solutions. The reader should refer to the literature for a discussion of the weighting process and a description of the various error contributions to be considered for G. Solving [3] relies on the singular value decomposition of the coefficient matrix A: A ¼ U S VT
½4
where U and V are orthogonal matrices of dimension n n and m m, respectively, and S is a diagonal matrix with the same dimension n m as A. The superscript T indicates transposition. The diagonal of S contains the nonnegative singular values si (i ¼ 1, y, n) sorted in decreasing size. The number of singular values greater than zero, p, is equal to the rank of A and reveals the actual number of independent equations in [3]. All subsequent singular values sp þ 1, y, sn are zero, and A can be rewritten as A ¼ Up Sp VpT
½5
where Up and Vp are trimmed versions of U and V of dimension n p and m p, respectively, containing only the first p columns of U and V. Sp is the p p submatrix of S consisting of the first p rows and columns only. With these quantities, the smallest least-squares solution bˆ of [3], the covariance matrix of the solution covðbˆÞ, and the resolution matrix R are given by T bˆ ¼ Vp S1 p Up G
½6
T covðbˆÞ ¼ Vp S2 p Vp
½7
R ¼ Vp VpT
½8
In these expressions, the superscript ‘ 1’ indicates the inverse of the matrix, and superscript ‘ 2’
indicates the product of the inverse matrix with itself. The resolution matrix R describes the relationship between the optimal solution bˆ obtained from the underdetermined system [3] with the ‘true’ solution b that one would find if [3] contained sufficient information and was full rank: bˆ ¼ R b
½9
Every component bˆj of the solution of the underdetermined system [3] can thus be represented as a linear combination of the ‘true’ unknowns P bˆj ¼ k rjk bk involving coefficients from row j of R. The analysis of the resolution matrix R and the deciphering of the real significance of the calculated bˆj are important steps for a meaningful interpretation of the results of underdetermined systems. Only if the resolution matrix R is diagonal (this is the case if [3] is full rank) will the calculated unknowns represent the ‘true’ unknowns. The section inverse approach described above has been applied by Ganachaud and Wunsch using hydrographic, nutrient, and oxygen data from 20 WOCE sections worldwide. Results for section integrated top-to-bottom net nitrate fluxes and divergences are shown in Figure 4. Net nitrate transports are found to be southward throughout the Atlantic and Indian Oceans and northward in the South Pacific. However, for many sections the uncertainties are of similar magnitude as the flux values themselves, leaving the net transports essentially indistinguishable from zero. Nitrate divergences in many regions of the world ocean indicate nutrient sinks in the upper part of the water column and nutrient sources below, consistent with the concept of biological production in surface waters and subsequent particle remineralization and release of nutrients below.
Estimating Carbon Export Fluxes with the Adjoint Method The section inverse approach described above requires hydrographic, nutrient, and tracer data along all interfaces of the domains for which budget equations are formulated. It is assumed that tracer fields and flows are in steady state, precluding the use of time-dependent tracers, like the CFCs. The budget equations could be generalized to include the time rate of change of tracer inventory inside the domain. However, one would need repeated tracer measurements inside the domain and along all its boundary for all times considered in the model to correctly describe the temporal inventory changes and the
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INVERSE MODELING OF TRACERS AND NUTRIENTS
305
60° N NITA 1 × 102 kmol s−1
10
80
20 × 102 kmol s−1
30° N
24.3(67m)
Uncertainty
26.44(162m)
+1 mol yr−1 m−2 −50
−10
−1 mol yr−1 m−2
26.44(183m)
−70 0° −210
−410
−160
−180
270
−60
−240 26.8(378m)
−70
30° S
190
−90
−240
60° S
4340
4620
4090 4300
60° W
0°
60° E
120° E
4700
180° W
120° W
Figure 4 Global dissolved nitrate transports and divergences. The length of each arrow corresponds to the nitrate transports between continents. The open boxes behind each arrow indicate the uncertainty (one standard deviation). Between sections, nitrate divergences are indicated by the solid boxes, either top-to-bottom (single box) or surface/deep (double box). Adapted from Ganachaud A and Wunsch C (2002) Oceanic nutrient and oxygen transport and bounds on export production during the World Ocean Circulation Experiment. Global Biogeochemical Cycles 16 (doi:10.1029/2000GB001333), with permission from the American Geophysical Union.
tracer fluxes across the boundaries. These massive data requirements cannot be met for any known transient tracer, and it seems necessary to adapt the inverse methodology for inclusion of sparse timedependent and steady-state tracers. A hybrid model consisting of forward and inverse steps and utilizing the Lagrange multiplier method of constrained variational optimization for fitting the model to tracer data has been developed for this purpose. The model exploits data for many tracers, including nutrients, radiocarbons, and CFCs. The objective of the model is to find optimal threedimensional (3-D) global ocean flows, biological production rates, and depth-dependent downward particle fluxes that explain the observed tracer, nutrient, and oxygen distributions best. The optimization is done iteratively, varying the flows as well as
the biogeochemical parameters systematically until the agreement between model simulated tracer fields and observations is optimal. The particular model has a rectangular grid (Figure 5), where grid cell boundaries are not required to match lines of available data. The layout of the grid is decoupled from the that of the available data, and individual grid cells (boxes) may be void of any data. Model tracer values are defined at the center of the boxes, whereas flows are defined on the interfaces. Biological production of particulate material occurs in the top model layers representing the euphotic zone. Particle fluxes below the euphotic zone are assumed to decrease with depth following a functional relationship from the literature
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jP ðzÞ ¼ a ðz=zEZ Þb
½10
306
INVERSE MODELING OF TRACERS AND NUTRIENTS
Gas exchange, heat, and freshwater fluxes
ZEZ
Export production
wi Depth
uj
vi
Water-column remineralization
wj 3-D circulation
Particle flux
Bottom remineralization Sedimentation
Figure 5 Vertical model grid and definition of model parameters. Reproduced from Schlitzer R, Usbeck R, and Fiscjer H (2004) Inverse modeling of particulate organic carbon fluxes in the South Atlantic. In: Wefer G, Mulitza S, and Ratmeyer V (eds.) The South Atlantic in the Late Quaternary – Reconstruction of Material Budget and Current Systems, pp. 1–19. Berlin: Springer, with permission of Springer Science Business Media.
where a is the particle flux at the base of the euphotic zone, zEZ, and represents the export production. The parameter b determines the shape of particle flux profile and thus controls the depth of remineralization. Large values for b correspond to steep particle flux decreases and thus large remineralization rates just below the euphotic zone, whereas values for b close to zero result in almost constant particle fluxes with depth with little remineralization in the water column and most of the particle export reaching the ocean floor. Export production a and remineralization parameter b vary geographically, and goal of the model runs is to infer optimal values for a and b (in addition to flow velocities v) based on the available nutrient and tracer data. Figure 6 shows a schematic overview of the computational strategy. All quantities listed under model parameters are to be determined by the model. These parameters are combined in the vector of independent parameters p . They have to be initialized to start the calculation (see the literature for initialization strategies), and they will be varied in a systematic way in the course of the calculations. Using the initial independent parameters, the model simulates the distributions of all steady-state (temperature, salinity, oxygen, phosphate, nitrate, total inorganic carbon, alkalinity, and radiocarbon) and transient (CFCs) tracers, as in normal forward models. The simulated tracer concentrations mi are combined in a ˜ All dependent vector of dependent parameters p.
parameters in p˜ can be calculated uniquely from the independent parameters p using the tracer budget equations for the boxes of the model. Once calculated, the simulated tracer distributions can be compared with measurements. Traditionally, this model/data comparison is done subjectively by analyzing the misfit fields and hypothesizing possible causes for the misfits. Model flows or biogeochemical parameters are then modified, hoping that new simulations with the modified parameters lead to more realistic tracer fields. This manual tuning has been used successfully with small box models; however, for problems with many thousands of parameters, such as the one described above, it is impractical and not successful in most cases. The Lagrange method of constrained optimization (termed ‘adjoint method’ in meteorology and oceanography) offers an alternative to manual parameter tuning. This method varies and optimizes the independent parameters p automatically, while satisfying the set of model equations consisting of all tracer budgets for all boxes exactly. The model equations are usually represented in homogeneous form Ej ¼ 0. The adjoint method can be applied to very large problems with hundred thousands of parameters or more. Here, the evaluation of the model performance is done objectively using a suitably defined cost function F. A cost function typically contains terms for all unwanted features of the model, most importantly terms that penalize deviations between model simulated tracer values mi and
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INVERSE MODELING OF TRACERS AND NUTRIENTS
307
Initialize
Model parameters
Model fields
Temperature
Nitrate
Salinity
Silicate
Air−sea fluxes
Phosphate
∑CO2
Export production
Oxygen
Alkalinity
Particle fluxes
CFC
Radiocarbon
Circulation Simulation
Mixing coefficients
Adjoint model
Improve model parameters
Compare with observed fields
Optimum? Figure 6 Schematic overview of model calculations performed for every iteration of the optimization process.
observations di7si: mi d i 2 ˜ ¼?þ Fðp ; pÞ þ? si
½11
A large value of F indicates that the simulated tracer concentrations differ much from the observations. Bringing the model close to the data, therefore, is equivalent to minimizing F, subject to satisfying the model equations (tracer budget equations) Ej ¼ 0 exactly. Details on the Lagrange multiplier method used for the constrained minimization of F can be found in textbooks and the literature on data assimilation or constrained data fitting. Overall, this method allows calculating the direction in parameter space rFp of steepest decrease of F (the negative gradient of F with respect to p ), which can then be used by a descent algorithm to arrive at a new, modified vector of independent parameters p . It is guaranteed that new simulations using the modified p will lead to more realistic tracer simulations and a smaller value of the cost function F. This procedure is repeated until no significant further decrease in F can be achieved or a limit on the number of iterations is reached (see Figure 6). The adjoint step in the calculations replaces the subjective evaluation of misfits mentioned above and provides an automatic
‘learning’ step of the model based on the current model/data misfits. The computational cost of the adjoint run is comparable to the cost of the simulation. The number of iterations required to reach the minimum of F depends on the initial p , but is usually large. It is therefore important to use efficient implementations of the simulation and adjoint steps to keep the computation time of a single iteration as short as possible. Figure 7 shows the observed and simulated radiocarbon values in the bottom waters of the world ocean after optimization. Following common practice, 14C concentrations are given in D-notation expressing the per mille 14C concentration difference of a given sample from the 14C standard (wood grown in 1890, decay corrected to 1950). A sample with D14C ¼ 100%, for instance, has a 14C concentration 100% (or 10%) lower than that of the standard. In agreement with the data and consistent with the general concept of the global thermohaline circulation, the model yields highest D14C values in the North Atlantic ( 70% to 80%) and lowest radiocarbon concentrations in the northeast Pacific ( 235%). D14C values in the Southern and Indian Oceans are intermediate and range between 150 and 165%. There are clear signs of tongues with elevated D14C values along the major deep and bottom water spreading paths. This includes the
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(c) 2011 Elsevier Inc. All Rights Reserved.
30° S
200
60° S
150
180°W
0
20
90°W
10
0
0
0°
5 15
90°E
Ocean data view
Figure 7 Bottom water D14C simulated by the model (color-shaded field) and from data (colored dots). The mean model-data D14C difference in areas not affected by bomb-14C is only þ 1.3%, and the root-mean-square difference amounts to only 5.2%. From Schlitzer R (2006) Assimilation of radiocarbon and chlorofluorocarbon data to constrain deep and bottom water transports in the world ocean. Journal of Physical Oceanography 37: 259–276.
EQ
30° N
60° N
55
200
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INVERSE MODELING OF TRACERS AND NUTRIENTS
southward flow of North Atlantic Deep Water (NADW) in the North Atlantic and the northward spreading of Antarctic Bottom Water (AABW) in the western South Atlantic as well as in the South Pacific and the Indian Oceans. The quantitative agreement between simulated and measured D14C values is excellent, and the model-data difference is only 1.3% on average. The root-mean-square (rms) difference amounts to only 5.2% and is of the same order of magnitude as the uncertainties of the radiocarbon data. In the deep North Atlantic and parts of the Southern Ocean, the measurements are systematically higher than the model simulations, owing to the contribution of bomb-14C in these waters. Bombradiocarbon is not included in the model, and contaminated data values are not assimilated in the model. The model fluxes of particulate organic carbon (POC) at the base of the euphotic zone (carbon export production) necessary to reproduce the observed oxygen and nutrient fields are shown in Figure 8. The spatial patterns of carbon export resemble the patterns of primary production in published maps showing high fluxes in coastal upwelling areas off West Africa, along the West American coast, in the Arabian Sea and Bay of Bengal, in the northwest Atlantic and north and tropical Pacific, in the area of the Indonesian archipelago, and in the Southern Ocean. Highest values in the productive areas are on the order of 5–10 mol C m 2 yr 1; in the oligotrophic, open-ocean regions they amount to between 0.5 and 2 mol C m 2 yr 1. The export of POC in the model predominantly occurs at mid-latitudes and in the Southern Ocean, and the contribution of the Northern Hemisphere is comparatively small. Globally integrated, the POC export in the model amounts to about 10 Gt C yr 1. Error analysis of the solution vector p is possible but computationally very expensive. An eigenvector/ eigenvalue analysis of the inverse Hessian matrix H 1 (the Hessian is the square matrix of second partial derivatives of F with respect to parameters p ) reveals directions in parameter space that are well determined (eigenvectors associated with large eigenvalues; values of F increase rapidly when moving away from optimal point along the direction of the eigenvector) and directions that are only poorly determined by the model (eigenvectors associated with smallest eigenvalues; values of F increase slowly when moving away from optimal point along the direction of the eigenvector). The ratio of largest and smallest eigenvalues of H 1 is a measure of the anisotropy of F around the optimal solution p . For relatively small problems with a few hundred parameters in p , the Hessian matrix, its inverse, and the associated eigenvalues and
eigenvectors can actually be calculated and a rigorous error analysis of the solution is possible. For large problems with hundred thousands of parameters, this will remain impossible for the foreseeable future. Strategies for obtaining sensitivity information at least for some directions in parameter space or for finding the largest eigenvalues and associated eigenvectors are described in the literature.
Conclusion Recent progress in applications of inverse models to problems in physical and biogeochemical oceanography clearly shows that inverse methodology is well advanced and being used by a growing number of researchers. Mathematical techniques exist that exploit the available data better than before and produce new and important scientific results. These methods successfully cope with problems, such as sparseness of data and incompleteness of information. The advances were possible because of the tremendous technological progress in computer hardware, combined with breakthroughs in the development of efficient and innovative algorithms. These algorithms finally allowed the numerical application of mathematical principles, such as the Lagrange multiplier method, whose theory was established for centuries already. Still, the widespread use of inverse methods would not have been possible, if there had not been at the same time an increased awareness among scientists for the need of integrated, global databases and an increased willingness to contribute to these data sets. Much more data will become available in the future, enabling inverse-type studies that are still impossible today. While ship observations will continue to be important, the new autonomous floats, gliders, or profiling instruments on moorings will provide data in near real time, even from remote and inaccessible regions. New satellite sensors will complement the water column data with high-resolution data from the ocean surface, and new geochemical programs, such as GEOTRACES, will produce highquality data of the ocean’s trace elements and isotopes, including micronutrients such as Fe and Zn, in unprecedented quality and coverage.
See also Inverse Models.
Further Reading Ganachaud A and Wunsch C (2002) Oceanic nutrient and oxygen transport and bounds on export production during the World Ocean Circulation Experiment.
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INVERSE MODELING OF TRACERS AND NUTRIENTS
Global Biogeochemical Cycles 16 (doi:10.1029/2000GB 001333). Gill PE, Murray W, and Wright MH (1981) Practical Optimization. London: Academic Press. Hestenes MR (1975) Optimization Theory. New York: Wiley. Kasibhatla P, Heimann M, Hartley D, Mahowald N, Prinn R, and Rayner P (eds.) (1998) AGU Geophysical Monograph Series, Vol. 114: Inverse Methods in Global Biogeochemical Cycles. Washington, DC: American Geophysical Union. Menke W (1984) Geophysical Data Analysis: Discrete Inverse Theory. San Diego, CA: Academic Press. Rintoul S and Wunsch C (1991) Mass, heat, oxygen and nutrient fluxes and budgets in the north Atlantic Ocean. Deep Sea Research 38(supplement): S355--S377. Schlitzer R (2000) Applying the adjoint method for global biogeochemical modelling. In: Kasibhatla P, Heimann M, Rayner P, Mahowald N, Prinn RG, and Hartley DE (eds.) AGU Geophysical Monograph Series, Vol. 114: Inverse Methods in Global Biogeochemical Cycles, pp. 107--124. Washington, DC: American Geophysical Union. Schlitzer R (2002) Carbon export fluxes in the Southern Ocean: Results from inverse modeling and comparison with satellite based estimates. Deep Sea Research II 49: 1623--1644. Schlitzer R (2004) Export production in the Equatorial and North Pacific derived from dissolved oxygen, nutrient and carbon data. Journal of Oceanography 60: 53--62. Schlitzer R (2006) Assimilation of radiocarbon and chlorofluorocarbon data to constrain deep and bottom water transports in the world ocean. Journal of Physical Oceanography 37: 259--276.
311
Schlitzer R, Usbeck R, and Fiscjer H (2004) Inverse modeling of particulate organic carbon fluxes in the South Atlantic. In: Wefer G, Mulitza S, and Ratmeyer V (eds.) The South Atlantic in the Late Quaternary – Reconstruction of Material Budget and Current Systems, pp. 1--19. Berlin: Springer. Tarantola A (1983) Inverse Problem Theory: Methods for Data Fitting and Model Parameter Estimation. New York: Elsevier. Thacker WC and Long RB (1988) Fitting dynamics to data. Journal of Geophysical Research 93: 1227--1240. Wunsch C (1978) The North Atlantic general circulation west of 501 W determined by inverse methods. Reviews of Geophysics 16: 583--620. Wunsch C (1996) The Ocean Circulation Inverse Problem. Cambridge, UK: Cambridge University Press.
Relevant Websites http://cdiac.esd.ornl.gov – Carbon Dioxide Information Analysis Center. http://whpo.ucsd.edu – Clivar and Carbon Hydrographic Data Office. http://www.coriolis.eu.org – Coriolis. http://www.ewoce.org – eWOCE: Electronic Atlas of WOCE Data. http://www.nodc.noaa.gov – National Oceanographic Data Center, NOAA. http://odv.awi.de – Ocean Data View.
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INVERSE MODELS C. Wunsch, Massachusetts Institute of Technology, Cambridge, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1368–1374, & 2001, Elsevier Ltd.
Introduction
w
Inverse methods are formal procedures for making inferences about the ocean (or any other physical system) by using observations in combination with dynamical and kinematic models. As such, they are a part of the general methods of statistical inference done in the presence of known or assumed kinematic and dynamical constraints. Their first use in oceanography occurred in 1977 as a method for addressing the famous so-called level-of-no-motion problem in the oceanic general circulation, and the determination of the general circulation remains a major area of application. Subsequently, inverse methods became a central element of ocean acoustic tomography. More recently, they have begun to be applied widely in all areas of oceanography including biogeochemical problems. In mathematical usage, ‘inverse methods’ often describe procedures directed at solving a variety of ill-posed problems, in the absence of observational noise. Although the terminology is much the same, noise-free observations exist only in textbooks, and this literature is useful, but tangential, to the oceanographic problem. ‘Inverse methods’ are often used in conjunction with, and thereby confused with, ‘inverse models’. For historical reasons, and mathematical convention, one denotes many systems as being ‘forward’ or ‘direct’ problems or models. A simple example derives from a supposed theory that produces a rule for a variable, perhaps oceanic temperature, y, as a function of z, in the form y ð z Þ ¼ a 0 þ a1 z þ a 2 z 2 þ a3 z 3 ¼
3 X
ai z 3
½1
i¼0
If the theory tells us the ai, we can calculate y for any value of z. The theory is often labeled a ‘forward’ or ‘direct model’ and, more generally, may be either an analytical or a very complex numerical one. Equation [1] is called a ‘forward solution’. If, on the other hand, y(z) were known, but one or more of the ai were not, one would have an inverse model, also
312
given by equation [1], the label ‘inverse’ being employed only as a matter of convention – because there is a previously studied ‘forward’ version. Another example comes from the classical advection/diffusion/decay equation for a concentration tracer, C, @C @2C k 2 þ lC ¼ qðzÞ @z @z
½2
where q is a source, w is the vertical advective velocity, k is a diffusion coefficient and l is a decay constant. All are known constants or functions of z. Given suitable boundary conditions, e.g., C(z ¼ 0)C0, C(z-N) ¼ 0, one has a well-understood, well-posed problem, in which eqn [2] and its boundary conditions are labeled as ‘forward.’ But an equally compelling, and commonplace problem is the following: Given C(z) and l, what are w and k? This type of problem occupies much of the mathematical literature on inverse problems. The oceanographers’ version is, however, more likely to be: Given observations, yi ¼ Cðzi Þ þ ni
½3
where ni are noise, at a finite number of discrete positions zi, what are w, k? Many variations of this problem are possible; for example, given noisy or uncertain observations or estimates of C, w, k, l what was the boundary condition C0? In this context, equation [2], along with any other available information, would now be described as the ‘inverse model,’ with the formal knowns and unknowns of the forward problem having in part, been interchanged. Parameter determination from noisy observations goes back at least to Legendre and Gauss. The modern generalization, inverse theory, is often traced to the pioneering work of G. Backus and F. Gilbert in the solid-earth geophysics context. They initiated the study of model systems in which the number of formal unknowns greatly exceeds (perhaps infinitely so), the number of available data by exploiting the existence of an underlying differential or partial differential system. This form of inverse theory is rooted in functional analysis and is highly developed; see Parker (1994), or for oceanographic applications, Bennett (1992), in the Further Reading. In oceanographic practice, even the simplest models are almost always reduced to some numerical, discrete form, either by expansion into a finite set of
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INVERSE MODELS
modes or by finite differences or related methods. The polynomial form eqn [1] is already discrete, with four parameters. Many other forms are possible. For example eqn [2], when written in finite differences, becomes Ciþ1 Ci Ciþ1 2Ci þ Ci1 ki þ lCi ¼ qi ; ziþ1 zi ziþ1 2zi þ zi1 1rirM wi
½4
which is a set of M simultaneous linear equations in the finite discrete set wi, ki, l. Alternatively, if w, k are constant, an analytic solution, subject to a fixed surface concentration C(z ¼ 0) ¼ C0 is readily seen to be (
" #) w 4lk 1=2 ; 1þ 1þ 2 CðzÞ ¼ C0 exp z 2k w
zr0 ½5
yi ¼ Ci þ ni, the problem has become one of finding an estimate, x˜, satisfying, Ex þ n ¼ y
" #) w 4lk 1=2 1þ 1þ 2 Cðzi Þ ¼ C0 exp zi 2k w
Ciþ1 Ci xiþ1 xi B ziþ1 zi ziþ1 zi ½6
Because eqns [4] and [6] are algebraic equations in w, k, l the problem of determining them has been reduced from that of an infinite-dimensional Hilbert or Banach space to that of an ordinary finite-dimensional vector space. No matter how great the value of M, the corresponding mathematical simplifications render the methods of inverse theory much more transparent in this case. Reduction of the functional analysis methods of Backus and Gilbert to that of finite-dimensional spaces appears to begin with Wiggins, who used the singular value decomposition of Eckart and Young. In practice, almost all real inverse problems are solved on computers; they are thus automatically discrete and of finite dimension and this mathematical representation is the most useful one. Note that eqn [6] is nonlinear in the parameters, showing that the same problem can be rendered linear depending upon exactly how it is formulated.
Example Assuming then, that observations always have a noise component, we can proceed to estimate whatever is known. For the simple power law of eqn [1], let us suppose that the ‘truth’ is yðzÞ ¼ 30 0:0005z2 ;
100r z r0
½7
but the theory says that it could actually be of the general form (1) so that the correct answer would be a0 ¼ 30, a2 ¼ 5 104, ai ¼ 0, otherwise. Defining
½8
Here the matrix vector notation x ¼ [ai] (called the ‘state vector’), n ¼ [ni], y ¼ [yi] is being used, and E is the coefficient matrix. As stated, this is now a problem in polynomial regression theory, and much of inverse theory overlaps that branch of statistics. The tracer problem (eqn [2]) can be reduced to this same form. Suppose that wi, ki, are believed constant, independent of i, and that we have noisy measurements xi [3] of Ci. Then one can readily make the substitution Ci-xi in eqn [4], producing a set of simultaneous equations for w, ki, where the coefficient matrix E has elements depending upon
which can be evaluated at z¼ zi to produce (
313
½9
etc. and y now involves lxi, qi, etc. n is a noise-vector representing the errors introduced by the observational noise, and any misrepresentation of the true full physics by eqn [2]. In practice, the structure of n can be a complicated function of the structure of E, because equations such as (9) render the problem, in a rigorous sense, nonlinear, too. This nonlinearity is often ignored and there are many inverse problems where E is known exactly. Equation [8] can now be regarded as the ‘model’ instead of, or in addition to, eqn [2]. The state vector consists of the two unknowns xT ¼ [w, ki]T (superscript T denotes a transpose; all vectors here are column vectors). The form of [8] obviously generalizes to any number of unknown elements, N, in x. A common method for dealing with equations of this type is to seek a least-squares solution that renders the noise vector n as small as possible: J ¼ nT n ¼ ðy ExÞT ðy ExÞ
½10
By taking the derivatives of the ‘objective’ or ‘cost’ function J with respect to the elements of x and setting them to zero, one obtains the ‘normal equations.’ Here, one finds, using this matrix notation, that the so-called normal equations are ET Ex ¼ ET y
½11
1 x˜ ¼ ET E y
½12
whose solution is
and which permits an estimate of the noise unknowns, 1 ½13 n˜ ¼ y Ex˜ ¼ y E ET E y
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314
INVERSE MODELS
The forward model [1] was used to generate ‘data’ at seven observation points, and these values were then corrupted with noise having a standard deviation of 1. The least-squares solution produces a state vector x˜T ¼ [31, 0.02, 0.002, 6 105, 4 107], and the corresponding estimated y˜ is shown in Figure 1. It passes near the data points but is clearly not ‘correct’ in the sense of reproducing the known true values. Although a very easy-to-use and common parameterdetermining procedure, least-squares in this form is not an inverse method (statisticians call it ‘curve-fitting’). The reasons for seeking a more powerful method are easy to see. At least as written, one can raise a number of questions about the solution: 1. Why should the smallest mean-square noise, nTn, be regarded as the correct solution to choose? 2. What would happen if the inverse of ETE failed to exist? (E corresponding to [7] is a Vandermonde matrix, and known to be very badly conditioned.) 3. Suppose one knew that some of the noise values were likely to be much greater than others: Could that information be used? 4. Suppose one had a reasonable idea of the magnitude of the ai, i.e., of x: Cannot that information be used? 5. Just how reliable is the solution given the noisiness of the data, and the particular structure of the observations? 6. Are some observations more important than others in determining the solution? 7. Could one require k>0 in the advection–diffusion problem? 8. In general, there are M eqns in (8) and N elements of x. When M>N, the system appears to be comfortably overdetermined. But that comfort disappears
when one recognizes that N noise unknowns have also been determined in eqn [13], producing a total of M þ N previously unknown values. Is it still sensible to term the problem ‘overdetermined.’ 9. Problems such as the advection-diffusion one involve a quantity C that is necessarily positive and that often renders the noise statistics highly nonGaussian. How can one understand how that affects the best solution? These and other issues lead one to find other means to make an estimate of x. It is possible to modify the least-squares procedure so that it will produce solutions obtained by other methods; this correspondence has led, in the literature, to serious confusion about what is going on.
Two More Solution Methods There are a number of methods available for estimating the elements of x, n that produce useful answers to some or all of the questions listed above. We will briefly describe two such methods, leaving to the references listed in Further Reading (see Menke, 1989; Tarantola, 1987; Munk et al., 1995; Wunsch, 1996) both the details and discussion of other approaches. Gauss–Markov Method
We postulate some a priori knowledge of the elements of x, n in a common, statistical, form. In particular, we assume that /xS ¼ /nS ¼ 0, where the brackets denote the expected value, and that there is some knowledge of the second moments, S ¼ /xxT S; R ¼ /nnT S.
0
_ 20
_ 20
_ 40
_ 40
z
z
0
_ 60
_ 60
_ 80
_ 80
_ 100 24 (A)
26
28 θ
30
32
_ 100 24
26
(B)
28 θ
30
32
Figure 1 (A) A hypothetical temperature, y, profile with depth, z, drawn as a solid line. ‘Observations’ of that profile at discrete depths contaminated with unit noise variance errors are shown as open circles. Dashed line is the ordinary least-squares fit. (B) Same as in (A), except that the dashed line is the Gauss–Markov solution computed using a priori knowledge of the statistics of the noise and the solution itself. The fit is less structured, and a specific error estimate is available at every point, showing the true line to lie within the estimated uncertainty.
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INVERSE MODELS
The so-called Gauss–Markov method (sometimes known as the ‘stochastic inverse’) produces a solution that minimizes the variance about the true value, i.e., it is a statistical method that minimizes not the sum of squares but, individually, all terms of the form /ðx˜ i xi Þ2 S where xi is the true value, and x˜ i is the estimate made. The solution is 1 x˜ ¼ SET ESET þ R y
½14
One then computes the singular value decomposition E0 ¼ UKVT
All of the available information has been used. A solution for the simple example is shown in Figure 1, where S¼diag ([100, 11010, 1106, 11010, 11010]), R¼{dij}. One obtains, x˜ T ¼½3170:8;1:41010 71105 ;6:1104 75:1104 ;4:5106 79:1106 ;4:9108 76:9108
½16
that is, the correct answer now lies within two standard errors and, although it is not displayed here, the P matrix provides a full statement of the extent to which the noise elements in x˜ are correlated with each other (which can be of the utmost importance in many problems). Least-squares by Singular Value Decomposition
As noted, least-squares can be modified so that it is more fully capable of producing answers to the questions put above. There are at least two ways to do this. The more interesting one is that based upon the singular value decomposition (SVD) and the socalled Cholesky decomposition of the covariance matrices, S ¼ ST/2S1/2, R ¼ RT/2R1/2. One takes the original eqn [18] and employs the matrices S, R to rotate and stretch E, n, y, x into new vector spaces in which both observations and solution have uncorrelated structures: RT=2 EST=2 ST=2 x þ RT=2 n ¼ RT=2 y
½17
E0 x0 þ n0 ¼ y0
½18
or
where E0 ¼ RT=2 EST=2 ; x0 ¼ ST=2 x; n0 ¼ RT=2 n; y0 ¼ RT=2 y
½19
½20
where K is a diagonal matrix, and U, V are orthogonal matrices. Let there be K nonzero values on the diagonal of K, and let the columns of U, V be denoted ui, vi respectively. Then it can be shown in straightforward fashion that the uTi uj ¼ dij ; vTi vj ¼ dij and that the solution to [18] is
with solution uncertainty (error estimate) 1 P ¼ /ðxx Þðxx ÞT S ¼ SSET ESET þR ES ½15 ˜ ˜
315
x0 ¼
K X k¼1
vk
N X uTk y0 þ ak vk lk k¼Kþ1
½21
where the ak are completely arbitrary. The physical solution is obtained from x˜ ¼ ST=2 x˜ 0 . A major advantage of this solution (and the uncertainly matrix, P, can also be computed for it), is that it explicitly produces the solution in orthonormal structures, vi, in terms of orthonormal structures of the observations ui, and separates these elements from the second sum on the right of equations [21], which defines the so-called null space of the problem. The null space represents the structures possibly present in the true solution, about which the equations [18] carry no information. A null space is present too, in the Gauss–Markov solution, but the corresponding null space vectors are given coefficients ak so as best to reproduce the a priori values of S. The SVD leads naturally to a discussion of what is called ‘resolution’ both of data and of the solution, and provides complete information about the solution. These issues, with examples, are discussed in Wunsch (1996) (see Further Reading). A Hydrographic Example
The first use of inverse methods in oceanography was the hydrographic problem. In this classical problem, oceanographers are able to determine the density field, r(z) at two nearby ‘hydrographic stations’ where temperature and salinity were measured by a ship as functions of depth. By computing the horizontal differences in density as a function of depth, and invoking geostrophic balance, an estimate could be made of the velocity field, up to an unknown constant (with depth) of integration. Although mathematically trivial, the inability to determine the integration constant between pairs of stations plagued oceanography for 100 years, and led to the employment of the ad hoc assumption that the flow field vanished at some specific depth z0. This depth became known as the ‘level-of-no-motion.’ But by employing physical requirements such as mass and salt conservation (or any other constraint involving
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INVERSE MODELS
B (cm s )
35
_1
_1
B (cm s )
the absolute velocity), it is straightforward to reduce the problem to one of deducing the actual flow field at the level-of-no-motion (and which is thus better called the ‘reference-level’). These constraints are readily written in the form of eqn [8], and solved with error estimates, etc. An example of the practical application of this method is shown in Figure 2. In this calculation, constraints were written for mass and salt conservation in a triangular region bounded by the US coastline, and a pair of hydrographic sections in which the Gulf Stream flowed into the region across one section and out again in the other. Direct velocity measurements obtained from the ship permitted additional equations to be written for the unknown flow velocity at the reference level. The resulting estimated absolute flow at the reference level is shown, with error estimates, in the top panels. The lower panels show the total estimated absolute velocity. In general, any available information can be used to estimate the unknowns of the system as long as one has a plausible estimate of the error contained in the resulting constraint. For the hydrographic problem in particular, a number of variations on the constraints have been proposed, under the labels of
_ 15 42 0
140
45
‘beta-spiral,’ ‘Bernoulli-method,’ etc., which we must leave to the references listed in Further Reading. Extensions
Like eqn [6], many inverse problems are nonlinear. Tarantola (1987) provides some general background and specific oceanographic applications may be seen in Mercier (1989), Mercier et al. (1991), and Wunsch (1994). The use of inequality constraints leads to the general subject of mathematical programming, a part of the wide subject of optimization theory (see Arthnari and Dodge, 1981). The Gauss–Markov solution method and the SVD version of least-squares have a ready interpretation, as minimum variance estimates of the true field. If the fields are all normally distributed, the solutions are also maximum-likelihood estimates, a methodology that is readily extended to nonGaussian fields.
Time-dependent Problems As originally formulated by Backus and Gilbert, and as exploited in most of the oceanographic literature
35 _ 15
49
61 60 58 _ 5 0
100
_ 10
80
55
50 49
160 100
_ 10
40 20 10
1000
_5 5
_5
60
5
1000
20
2000
Pressure (db)
Pressure (db)
10 _5
5
10
_ 10
3000
2000 5
3000 _ 15 _5
4000 _ 20
20 km
4000
20 km
Figure 2 Example of the inversion for the reference level velocity in a triangular region bounded on two sides by hydrographic sections crossing the Gulf Stream, and on the third side by the US coastline. Velocity contours are in centimeters. They are the sum of the so-called thermal wind, which involves setting the velocity to zero at a reference depth. The actual flow at that depth (the ‘reference level velocity’) is shown in the top panels as estimated, with uncertainties, from the singular value decomposition solution, with K ¼ 30. (From Joyce et al., 1986.) The ‘columnar’ structure, which is so apparent here, first appeared in inverse solutions, and was greeted with disbelief by those who ‘knew’ that oceanic flows were ‘layered’ in form. Acceptance of this type of structure is now commonplace.
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INVERSE MODELS
to date, the problems have been essentially static, with time evolution not accounted for. If one has a system that evolves in time, the observations and state vector, x, will also be time evolving. If one simply writes down the relationship between data and a model in which the two are connected by a set of equations, linear or nonlinear, one sees immediately that, mathematically, the problem is identical to that posed by eqn [8], in which the time of the observation or of the model calculation is just a bookkeeping index. The difficulty that arises is purely a practical one: the potentially extremely large growth in the number of equations that must be dealt with over long time spans. With a sufficiently large and fast computer, the most efficient way to solve such inverse problems would be the straightforward application of the same methods developed for static problems: Write out all of the equations explicitly and solve them all at once. In oceanographic practice however, one rapidly outstrips the largest available computers and the static methods become impractical if used naively. Fortunately, time-evolving oceanographic models have a very special algebraic structure, which permits one to solve the corresponding normal equations [11] by methods not requiring storage of everything in the computer at one time. Numerous such special methods exist, and go by names such as sequential estimators (Kalman filter and various smoothers), adjoint equations (Pontryagin principle or method of Lagrange multipliers), Monte Carlo methods, etc. The generic terminology that has come into use is ‘data assimilation’ borrowed from meteorological forecasting terminology. Such methods are highly developed, if often abused or misunderstood, and require a separate discussion. What is important to know, however, is that they are simply algorithmically efficient solutions to the inverse problem as described here.
Common Misconceptions and Difficulties Some chronic misunderstandings and difficulties arise. The most pernicious of these is the attempt to use inverse models, which are physically inconsistent with the known forward model or physics. This blunder corresponds, for example, in the simplest model above, to imposing a linear relationship (model) between y and z, when it is clear or suspected that higher powers of z are likely to be present. Some writers have gone so far as to show that such an inversion does not reproduce a known forward solution, and then declared inverse methods to be failures. Another blunder
317
is to confuse the inability to resolve or determine a parameter of interest, when the data are inadequate for the purpose, with a methodological or model failure. Inverse methods are very powerful tools. Like any powerful tool (a chain saw, for example), when properly used they are useful and even essential; when improperly used they are a grave danger to the user.
See also Data Assimilation in Models. Elemental Distribution: Overview. General Circulation Models. Tomography. Tracer Release Experiments.
Further Reading Arthnari TS and Dodge Y (1981) Mathematical Programming in Statistics. New York: Wiley. Bennett AF (1992) Inverse Methods in Physical Oceanography. Cambridge: Cambridge University Press. Eckart C and Young G (1939) A principal axis transformation for non-Hermitian matrices. Bulletin of the American Mathematical Society 45: 118--121. Joyce TM, Wunsch C, and Pierce SD (1986) Synoptic Gulf Stream velocity profiles through simultaneous inversion of hydrographic and acoustic doppler data. Journal of Geophysical Research 91: 7573--7585. Lanczos C (1961) Linear Differential Operators. Princeton, NJ: Van Nostrand. Liebelt PB (1967) An Introduction to Optimal Estimation. Reading, MA: Addison Wesley. Menke W (1989) Geophysical Data Analysis: Discrete Inverse Theory 2nd edn. New York: Academic Press. Mercier H (1989) A study of the time-averaged circulation in the western North Atlantic by simultaneous nonlinear inversion of hydrographic and current-meter data. DeepSea Research 36: 297--313. Mercier H, Ollitrault M, and Le Traon PY (1991) An inverse model of the North Atlantic general circulation using Lagrangian float data. Journal of Physical Oceanography 23: 689--715. Munk W, Worcester P, and Wunsch C (1995) Ocean Acoustic Tomography. Cambridge: Cambridge University Press. Parker RL (1994) Geophysical Inverse Theory. Princeton, NJ: Princeton University Press. Seber GAF (1977) Linear Regression Analysis. New York: Wiley. Tarantola A (1987) Inverse Problem Theory. Methods for Data Fitting and Model Parameter Estimation. Amsterdam: Elsevier. Van Trees HL (1968) Detection, Estimation and Modulation Theory. Part I. New York: Wiley. Wiggins RA (1972) The general linear inverse problem: implication of surface waves and free oscillations for earth structure. Reviews in Geophysics and Space Physics 10: 251--285.
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Wunsch C (1977) Determining the general circulation of the oceans: a preliminary discussion. Science 196: 871--875. Wunsch C (1994) Dynamically consistent hydrography and absolute velocity in the eastern North
Atlantic Ocean. Journal of Geophysical Research 99: 14071--14090. Wunsch C (1996) The Ocean Circulation Inverse Problem. Cambridge: Cambridge University Press.
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IR RADIOMETERS C. J. Donlon, Space Applications Institute, Ispra, Italy
principles described are applicable to satellite sensors treated elsewhere in this volume.
Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1374–1385, & 2001, Elsevier Ltd.
Introduction Measurements of sea surface temperature (SST) are most important for the investigation of the processes underlying heat and gas exchange across the air–sea interface, the surface energy balance, and the general circulation of both the atmosphere and the oceans. Complementing traditional subsurface contact temperature measurements, there is a wide variety of infrared radiometers, spectroradiometers, and thermal imaging systems that can be used to determine the SST by measuring thermal emissions from the sea surface. However, the SST determined from thermal emission can be significantly different from the subsurface temperature (471 K) because the heat flux passing through the air–sea interface typically results in a strong temperature gradient. Radiometer systems deployed on satellite platforms provide daily global maps of SSST (sea surface temperature) at high spatial resolution (B1 km) whereas those deployed from ships and aircraft provide data at small spatial scales of centimeters to meters. In particular, the development of satellite radiometer systems providing a truly synoptic view of surface ocean thermal features has been pivotal in the description and understanding of the global oceans. This article reviews the infrared properties of water and some of the instruments developed to measure thermal emission from the sea surface. It focuses on in situ radiometers although the general
Infrared Measurement Theory Infrared (IR) radiation is heat energy that is emitted from all objects that have a temperature above 0 K ( 273.161C). It includes all wavelengths of the electromagnetic spectrum between 0.75 mm and B100 mm (Figure 1) and has the same optical properties as visible light, being capable of reflection, refraction, and forming interference patterns. The following total quantities, conventional symbols and units provide the theoretical foundation for the measurement of IR radiation and are schematically shown in Figure 2. Spectral quantities can be represented by restricting each to a specific waveband.
• • •
• •
Radiant energy Q, is the total energy radiated from a point source in all directions in units of joules (J). Radiant flux f ¼ dQ/dt is the flux of all energy radiated in all directions from a point source in units of watts (W). Emittance M ¼ df/dA is the radiant flux density from a surface area A in units of W m2. This is an integrated flux (i.e., independent of direction) and will therefore vary with orientation relative to a nonuniform source. Radiant intensity I ¼ df/do is the radiant flux of a point source per solid angle o (steradian, sr) and is a directional flux in units of W sr1. Radiance L ¼ dI/d(A cos y) is the radiant intensity of an extended source per unit solid angle in a given direction y, per unit area of the source projected in the same y. It has units of W sr1 m2.
Figure 1 Schematic diagram of the electromagnetic spectrum showing the location and interval of the infrared waveband.
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IR RADIOMETERS
y (surface normal) Radiant flux
Surface area, A
x
Figure 3 Black-body radiance as a function of wavelength computed using eqns [1] and [2] for different target temperatures. z (A) y Area = A Radiant flux
Solid angle 2 = A /R
where h ¼ 6.626 1034 J s is Planck’s constant, c ¼ 2.998 108 m s1 is the speed of light and k ¼ 1.381 1023 J K1 is Boltzmann’s constant. The sea surface is considered a Lambertian source (i.e., uniform radiance in all directions) so that the spectral radiance Ll is related to Ml by Ll ¼
R x
z (B)
y (surface normal) d
Surface area, A
Radiant flux
Direction, x Projected source area, dA cos
z (C)
Figure 2 Schematic definition of (A) Emittance, E; (B) radiant intensity, I; (C) radiance, L.
Planck’s law describes the emittance of a perfectly emitting surface (or black body) at a temperature T in Kelvin. It is the radiant flux (f) per unit bandwidth centered at wavelength l leaving a unit area of surface in any direction in units of W m2 m1. Ml;T ¼
2phc2 l5 ðehc=lkT 1Þ
Ml p
½2
Figure 3 shows Ll computed for several temperatures as a function of wavelength. Considering temperatures of 273–310 K as representative of the global ocean, maximum emission occurs at a wavelength of 9.3–10.7 mm. Atmospheric attenuation is minimal at B3.5 mm, 9.0 mm and 11.0 mm that are the spectral intervals often termed atmospheric ‘windows’. Instruments operating within these intervals are optimal for sea surface measurements – especially in the case of satellite deployment where atmospheric attenuation can be significant. In the 3–5 mm spectral region. Ll is a strong function of temperature (Figure 3) highlighting the possibility to increase in radiometer sensitivity by utilizing this spectral interval. By measuring Ll using eqn [2] and inverting eqn [1], the spectral brightness temperature, B(T,l), rather than the temperature is determined because these equations assume that sea water is a perfect emitter or black body. In practice, the sea surface does not behave as a black body (it is slightly reflective in the infrared) and therefore its spectral and geometric properties need to be considered. The emittance of a perfect emitter at the actual temperature T, wavelength l, and view angle y is given by
½1
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MðT;l;yÞ ¼
pLðT;l;yÞ eðl;yÞ
½3
IR RADIOMETERS
B11 mm (rE0.0015) and following Kirchoff’s law
where the emissivity, eðl;yÞ , can be calculated using eðl;yÞ ¼
rl þ tl þ el ¼ 1
MðT;l;yÞ measured MðT;l;yÞ blackbody
321
½5
½4
which has a strong dependence on wavelength and viewing geometry. The effective emissivity, e, integrates eðl;yÞ over all wavelengths of interest for radiometer view angle y. Figure 4A shows the calculated normal reflectivity, r, of pure water as a function of wavelength for the spectral region 1–100 mm. Pure water differs only slightly from sea water in this context. Note that r is minimal at a wavelength of
where rl is the spectral reflectivity and tl is the spectral transmissivity. The e-folding penetration depth (i.e., the depth of 63% emission) or optical depth at a typical wavelength of 11 mm is E 10 mm, tl can be neglected and el can be calculated using el ¼ 1 rl
½6
Although the actual optical depth is wavelength dependent, it is clear that IR radiometers determine
(zenith angle = 40˚)
0.100
0.010
(h+v) h v 0.001
10
100 Wavelength (μm)
(A)
Sea surface emissivity ( = 1 _ )
1.00
0.95
0.90
0.85
0.80 (B)
0˚ 40˚ 50˚ 60˚ 70˚
10
100 Wavelength (μm)
Figure 4 (A) The normal reflectivity, r, of the pure water as a function of wavelength (full line). Also shown are the horizontal polarized (rh) and the vertically polarized (rv) components of r. (B) The spectral emissivity of pure water as a function of viewing zenith angle. h, height above sea surface.
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Figure 5 Schematic diagram showing the radiance components measured by an IR radiometer viewing the sea surface. Fov, Field of view.
the temperature of a very thin ‘skin’ layer of the ocean. This temperature is termed the sea surface skin temperature (SSST) and is distinct (although related) to the subsurface SST. Note that this is in contrast to the situation for short-wave solar radiation (having wavelengths of B0.4–0.7 mm) which, for clear water, penetrates to a depth of B100 m. Figure 4B shows the el of pure water computed from eqn [6] as a function of both viewing zenith angle and spectral wavelength. Inspection of Figure 4 reveals that the best radiant temperature measurement will be made when viewing a calm sea surface at an angle of 0–401 from nadir. An IR radiometer is an optical instrument designed to measure Ll entering an instrument aperture (Figure 5). The radiance measured by radiometer, LðT;l;yÞ , having a spectral bandwidth l, viewing the sea surface at a zenith angle y and temperature T is given by: LðT;l;y ¼
ða
xl ½eðl;yÞ B Tsurf ; l þ 1 eðl;yÞ BðTatm ; lÞ
0
surface because diffuse downwelling sky radiance measured by a radiometer after reflection at the sea surface is polarized.
IR Radiometer Design There are four fundamental components to all IR radiometer instruments described below. Detector and Electronics System to Measure Radiance and Control the Radiometer
A detector system provides an output proportional to the target radiance incident on the detector. There are two main types of detector: thermal detectors that respond to direct heating and quantum detectors that respond to a photon flux. In general, thermal detectors have a response that is weakly dependent on wavelength and can be operated at ambient temperatures whereas rapid response quantum detectors require cooling and are wavelength dependent.
þLpathðh;l;yÞ dl ½7 where xl is the spectral response of the radiometer, BðTsurf ; l and BðTatm ; l are the Planck function for surface temperature Tsurf, and atmospheric temperature Tatm, and Lpathðh;l;yÞ is the radiance emitted by the atmosphere between the radiometer at height h above the sea surface reflected into the radiometer field-of-view (FoV) at the sea surface. Note that the horizontal and vertical polarization components of reflectivity shown in Figure 4 are unequal. It is important to consider the polarization of surface reflectance when making measurements of the sea
Fore-optics System to Filter, Direct and Focus Radiance
All optical components have an impact on radiometer reliability and accuracy. Mirrors should be free of aberration to minimize unwanted stray radiance reaching the detector. Several materials have good reflection characteristics in the IR including, gold, polished aluminum, and cadmium. Care should be exercised when choosing an appropriate mirror substrate and reflection coating to avoid decay in the marine atmosphere. A glass substrate having a ‘hard’ scratch-resistant polished gold surface provides 498% reflectance and good environmental wear.
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IR RADIOMETERS
323
1.0 Detector transmission Window transmission Total transmission
Normalized transmission
0.8
0.6
0.4
0.2
0.0 6
8
10 12 Waveband (μm)
14
16
Figure 6 Normalized spectral transmission of an IR window and detector shown together with the combined total response.
Spectral filter windows and lenses require spectral properties that, together with the detector characteristics, define the overall spectral characteristics of a radiometer. Figure 6 shows the combined spectral response for a broadband radiometer together with the component spectral response of the window and detector. Environmental System to Protect and Thermally Stabilize the Radiometer
For any optical instrument intended for use in the harsh marine environment, adequate environmental protection is critical. Rain, sea-water spray, and high humidity can destroy a poorly protected instrument rapidly and components such as electrical connections and fore-optics should be resistant to these effects. In situ radiometer windows are particularly important in this context. They should not significantly reduce the incoming signal or render it noisy, and be strong enough to resist mechanical, thermal and chemical degradation. Figure 7 describes several common materials. Germanium (Ge) windows have good transmission characteristics but are very brittle. Sodium chloride (NaCl) is a low-cost, low-absorption material but is of little use in the marine environment because it is water-soluble. Zinc selenide (ZnSe) has high transmission, is nonhygroscopic
Figure 7 Spectral transmission for common IR window materials. (A) Germanium, (B) sodium chloride, (C) zinc selenide.
and resistant to thermal shock but is soft and requires a protective ‘hard’ surface finish coating. Certain window materials achieve better performance when antireflection (AR) coatings are used to minimize reflection from the window. For example, when an AR coating is used on a ZnSe window the transmission increases from B70% to B90%. Other coatings provide windows that polarize the incident radiance signal such as optically thin interference coatings and wire grid diffraction polarizers. Deposition of marine NaCl on all optical components (especially calibration targets) presents an unavoidable problem. Although NaCl itself has good infrared transmission properties (Figure 7), contaminated surfaces may become decoupled from temperature sensors and the noise introduced by the optical system will increase. Finally, adequate thermal control using reflective paint together with substantial instrument mass is required so that instruments are not sensitive to thermal shock. In higher latitudes, it may be necessary to provide an extensive antifreeze capability. Calibration System to Quantify the Radiometer Output
The role of a calibration system is to quantify the instrument output in terms of the measured radiance incident on the detector. Calibration techniques are specific to the particular design of radiometer and vary considerably from simple bias corrections to
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systems providing automatic precision two-point blackbody calibrations. Proper calibration accounts for the following primary sources of error:
• • •
the effect of fore-optics; unavoidable drifts in detector gain and bias; long-term degradation of components.
Finally, careful radiometer design and configuration can avoid many measurement errors. Examples include: poor focusing and optical alignment; filters having transmission above or below the stated spectral bandwidth (termed ‘leaks’); inadequate protection against thermal shock, stability, and reliability of the calibration system; poor electronics and the general decay of optomechanical components.
Application of IR Radiometers There are many different radiometer designs and deployment strategies ranging from simple singlechannel hand-held devices to complex spectroradiometers. Instrument accuracy, sensitivity and stability depends on both the deployment scenario and radiometer design. Accordingly, the following sections describe several different examples. Broadband Pyrgeometers
A pyrgeometer is an integrating hemispheric radiometer which, by definition, measures the bandlimited spectral emittance, E, so that the angle dependency in eqn [7] is redundant. They are used to determine the long-wave heat flux at the air–sea interface by measuring the difference between atmospheric and sea surface radiance either using two individual sensors (Figure 8A) or as a single combined sensor (Figure 8B). A thermopile detector is often used which is a collection of thermocouple detectors composed of two dissimilar metallic conductors connected together at two ‘junctions.’ The measurement junction is warmed by incident
radiance relative to a stable reference junction and a mV signal is produced. The response of a thermopile has little wavelength dependence and a hemispheric dome having a filter (B3–50 mm) deposited on its inner surface typically defines the spectral response. Direct compensation for thermal drift using a temperature sensor located at the thermopile reference junction is sometimes used but regular calibration using an independent laboratory blackbody is mandatory. An accuracy of o10 W m2 is possible after significant correction for instrument temperature drift and stray radiance contribution using additional on-board temperature sensors. Narrow Beam Filter Radiometers
Narrow beam filter radiometers are often used to determine the SSST for air–sea interaction studies and for the validation of satellite derived SSST and there are several low-cost instruments that provide suitable accuracy and spectral characteristics. Many of these use a simple thermopile or thermistor detector together with a low-cost broadband-focusing lens. Typically, they have simple self-calibration techniques based on the temperature of the instrument and/or detector. Consequently they have poor resistance to thermal shock and have fore-optics that readily degrade in the marine atmosphere. However, handled with care, these devices are accurate to 70.1 K, albeit with limited sensitivity. Precision narrow beam filter radiometers often use pyroelectric detectors that produce a small electrical current in response to changes in detector temperature forced by incident radiation. They have a fast response at ambient temperatures but require a modulated signal to operate. Modulation is accomplished by using an optical chopper having high reflectivity ‘vanes’ to alternately view a reference radiance source by reflection and a free path to the target radiance. The most common chopper systems are rotary systems driven by a small electric motor phase locked to the
Figure 8 (A) A typical design of a long-wave pyrgeometer. (B) A net-radiation pyrgeometer for determination of the net long-wave flux at the sea surface.
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IR RADIOMETERS
325
Figure 9 (A) Schematic layout of a rotary chopper; (B) schematic layout of a tuning fork chopper; (a) open; (b) closed.
detector output by an optoelectronic sensor Figure 9A. An alternative design driven by a small oscillating electromagnetic coil called a tuning fork chopper is shown in Figure 9B. As the coil resonates, the reflective vanes of the chopper oscillate alternately opening and closing an aperture ‘gap.’ Dynamic detector bias compensation is inherent when using an optical chopper. The detector alternately measures radiance from the sea surface Lsrc and a reference blackbody, Lbb (sometimes this is the detector itself) reflected by the chopper vanes resulting in two signals S1 ¼ Lbb þ d
½8
S2 ¼ Lsrc þ d
½9
Assuming Lbb remains constant during a short chopping cycle, the bias term d in eqns [8] and [9] is eliminated DS ¼ S1 S2 ¼ Lbb Lsrc
½10
It is important to recognize the advantages to this technique, which is widely used:
• • •
There is minimal thermal drift of the detector; The detector is dynamically compensated for thermal shock; A precise modulated signal is generated well suited to selective filtering providing excellent noise suppression and signal stability.
However, in order to compensate for instrument gain changes, an additional blackbody sources(s) is required. These are periodically viewed by the detector to provide a mechanism for absolute calibration. Either the black body is moved into the detector FoV or an adjustable mirror reflects radiance from the black body on to the detector.
Calibration cycles should be made at regular intervals so that gain changes can be accurately monitored and calibration sources need to be viewed using the same optical path as that used to view the sea surface. A basic ‘black-body’ calibration strategy uses an external bath of sea water as a high e (40.95) reference as shown in Figure 10. In this scheme, the radiometer periodically views the water bath that is stirred vigorously to prevent the development of a thermal skin temperature deviation. The view geometry for the water bath and the sea surface are assumed to be identical and, by measuring the temperature of the water bath the radiometer can be absolutely calibrated. An advantage of this technique is that eðl;yÞ is not required to determine the SSST. However, in practice, it is difficult to continuously operate a water bath at sea and surface roughness differences between the bath and sea surface are ignored. On reflection at the sea surface, diffuse sky radiance is polarized and, at Brewster’s angle (B501 from nadir at a wavelength of 11 mm), the vertical vpolarization is negligible for a given wavelength (Figure 11). Only the horizontal h-polarization component remains so that if the radiometer filter response is v-polarized (i.e., only passes v-polarized radiance), negligible reflected sky radiance is measured by the radiometer. In practice, because Brewster’s angle is very sensitive to the geometry of a particular deployment (approximately721) this technique is only applicable to deployments from fixed platforms and when the sea surface is relatively calm. Further, the use of a polarizing filter will significantly reduce the signal falling on the detector increasing the signal-to-noise ratio. The use of fabricated black-body cavities (Figure 12A) provides an accurate, versatile and, compact calibration system. Normally, two
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IR RADIOMETERS
Radiometer
Radiometer beam
Water overflow
Inner water bath
Water bath moved out from radiometer FoV which no views the sea surface
Gimball mounting
Overflow drain
Sea-water input
(A)
(B)
Figure 10 Schematic diagram showing the stirred water bath calibration scheme. (A) The radiometer in calibration mode, (B) the radiometer viewing the sea surface after the water bath has been moved out of the field of view (FoV).
black-body cavities are used, one of which follows the ambient temperature of the instrument and a second is heated to a nominal temperature above this. High e (40.99) is attained by a combination of specialized surface finish and black-body geometry. The cavity radiance is determined as a function of the black body temperature that is easily measured. Figure 12B shows a schematic outline of a typical black-body radiometer design using a rotary chopper
and Figure 12C provides a schematic diagram of a typical output signal. Note that for all calibration schemes, larger errors are expected beyond the calibrated temperature range which can be a problem for sky radiance measurements where clear sky temperatures of o200 K are common.
Multichannel Radiometers at = 11μm for angles 0 _ 90˚
1.0000
at = 11 μm
0.1000
0.0100
0.0010
h v 0.0001
0
20
40 60 Zenith angle (˚)
80
Figure 11 Polarization of sea surface reflection at 11 mm as a function of view angle. Total polarization is shown as a solid line.
The terms in eqn [7] are directly influenced by the height, h, of the radiometer above the sea surface and are different in magnitude for in situ and spacecraft deployments. In the case of a sea surface in situ radiometer deployment, Lpathðh;l;yÞ is typically neglected because h is normally o10 m unless the atmosphere has a heavy water vapor loading (e.g., 490%) or an aircraft deployment is considered. However, for a spacecraft deployment, this is a significant term requiring explicit correction. Conversely, the BðTatm ; lÞ term is critical to the accuracy of an in situ radiometer deployment but of little impact (except perhaps at the edge of clouds) for a satellite instrument deployment because Lpathðh;l;yÞ dominates the signal. A multispectral capability can be used to explicitly account for Lpathðh;l;yÞ in eqn [7] because of unequal atmospheric attenuation for different spectral wavebands. Multichannel radiometers are exclusively used on satellite platforms for this reason. Many in situ multichannel radiometers are designed primarily for the radiant calibration or validation of specific satellite radiometers and the development of satellite radiometer atmospheric correction algorithms. They often have several selectable filters matched to those of the satellite sensor.
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Figure 12 (A) A section through a black-body calibration cavity of the re-entrant cone design. (B) A typical black-body calibration radiometer using a rotary optical chopper. (C) A schematic diagram of a typical detector output signal, showing sea, reference, and calibration signals.
It is worth noting that multiangle view radiometers are also capable of providing an explicit correction for atmospheric attenuation. Often operated from satellite and aircraft, these instruments provide a direct measure of atmospheric attenuation by making two views of the same sea surface area at different angles using a geometry that doubles the atmospheric pathlength (Figure 13). The assumption is made that atmospheric and oceanic conditions are stationary in the time between each measurement.
or satellite
Spectroradiometers
A recent development is the use of Fourier transform infrared spectrometers (FTIR) that are capable of accurate (B0.05 K). High spectral resolution
1+dt
Figure 13 Schematic diagram of dual view, atmospheric path radiometer deployment geometry.
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(B0.5 cm1) measurements over a broad spectral range (typically B3–18 mm) as shown in Figure 14. The FTIR provides a unique tool for the development of new IR measurement techniques and investigation of processes at the air–sea interface. For example, the Marine-Atmospheric Emitted Radiance Interferometer (M-AERI) has pioneered a SSST algorithm that uses a narrow spectral region centered at 7.7 mm that is less susceptible to the influence of cloud cover and sky emissions that at 10–12 mm. The FTIR can also be used to measure air temperatures by viewing the atmosphere at B15 mm (a spectral region opaque due to CO2 emission) that are accurate to o0.1 K. Of considerable interest is the ability of an FTIR provide an indirect estimate of rðl;yÞ so that by using eqn [5], eðl;yÞ can be computed. The sky radiance spectrum has particular structures associated with atmospheric emission–absorbance lines (Figure 14A) that are physically uncorrelated with the smooth spectrum of rðl;yÞ (Figure 4). The spectrum of rðl;yÞ can be derived by subtracting a scaled BðTarm ; lÞ
spectrum to minimize the band-limited variance of the BðTsea ; lÞ spectrum (Figure 14B). Finally, direct measurement of the thermal gradient at the air–sea interface to obtain the net heat flux has been demonstrated using an FTIR in the laboratory. The FTIR uses the 3.3–4.1 mm spectral interval that has an effective optical depth (EOD) depending on the wavelength (EOD ¼ 0 mm at 3.3 mm whereas at 3.8 mm EOD ¼ 65 mm) demonstrating the versatility of the FTIR. However, measurement integration times are long and further progress is required before this technique is applicable for normal field operations. Thermal Imagers
Another recent development is the application of IR imagers and thermal cameras for high-resolution process studies such as fine-scale variability of SSST, wave breaking (Figure 15), and understanding air– sea gas and heat transfer. They are also used during
Figure 14 Spectra of emitted sky and sea view radiation measured by the M-AERI FTIR in the tropical Western Pacific Ocean on March 24, 1996. (A) Spectrum of sky radiance and (B) spectrum of corresponding sea radiance Sky measurements were made at 451 and zenith (red) above the horizon and ocean measurements were made at 451 below the horizon. The cold temperatures in the sky spectra show where the atmosphere is relatively transparent. The ‘noise’ in the 5.5–7 mm range is caused by the atmosphere being so opaque that the radiometer does not ‘see’ clearly the instrument internal black-body targets and calibration is void. The spectrum of upwelling radiation (B) consists of emission from the sea surface, reflected sky emission and emission from the atmospheric pathlength between the sea surface and the radiometer.
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Figure 15 A thermal image of a breaking wave. Each pixel is B25 25 cm and the wavelength of the camera is 8–12 mm. (Courtesy of D. Woolf.)
air–sea rescue operations providing a nighttime capability for detecting warm objects such as a life raft or survivors. These instruments use a focal plane array (FPA) detector (a matrix of individual detectors, e.g., 256 256) located at the focal point of incoming radiance together with a charge couple device that is used to ‘read’ the FPA. This type of ‘staring array’ system generates a two-dimensional image either by using a mechanical scanning system (in the case of a small FPA array) or as an instantaneous image. Larger FPA arrays are much more power efficient, lighter and smaller than more elaborate mechanical scanning systems. Rapid image acquisition (415 frames s1) is typical of these instruments that are available in a wide variety of spectral configurations and a typical accuracy of B70.1 K. However, considerable problems are encountered when obtaining sky radiance data due to the difficulty of geometrically matching sea and sky radiance data. The major problem with FPA detector technology is nonuniformity between FPA detector elements and drifts in detector gain and bias. Many innovative self-calibration methods which range in quality are used to correct for these problems. For example, a small heated calibration plate assumed to be at an isothermal temperature is periodically viewed by the detector to provide an absolute calibration. However, further development of this technology will
eventually provide extremely versatile instrumentation for the investigation of fine-scale sea surface emission.
Future Direction and Conclusions In the last 10 years, considerable progress has been made in the development and application of IR sensors to study the air–sea interface. The continued development and use of FTIR sensors will provide the capability to accurately investigate the spectral characteristics of the sea surface in order to optimize the spectral intervals used by space sensors to determine SSST. It can be expected that in the near future, new algorithms will emerge for the direct measurement of the air–sea heat flux using multispectral sounding techniques and the accurate in situ determination of sea surface emissivity. Although still in their infancy, the development and use of thermal cameras will provide valuable insight into the fine-resolution two-dimensional spatial and temporal variability of the ocean surface. These data will be useful in developing and understanding the sampling limitations of large footprint satellite sensors and in the refinement of validation protocols. Finally, as satellite radiometers are now providing consistent and accurate observations of the SSST (e.g., ATSR), there is a need for autonomous
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operational in situ radiometer systems for ongoing validation of their data. Such intelligent systems that are extremely robust against the harsh realities of the marine environment are currently being developed.
See also Air–Sea Gas Exchange. Heat and Momentum Fluxes at the Sea Surface. Radiative Transfer in the Ocean. Satellite Remote Sensing of Sea Surface Temperatures.
Further Reading Bertie JE and Lan ZD (1996) Infrared intensities of liquids: the intensity of the OH stretching band revisited, and the best current values of the optical constants H2O (1) at 251C between 15,000 and 1 cm1. Applied Spectroscopy 50: 1047--1057. Donlon CJ, Keogh SJ, Baldwin DJ, et al. (1998) Solid state measurements of sea surface skin temperature. Journal of Atmospherical Oceanic Technology 15: 775--787. Donlon CJ and Nightingale TJ (2000) The effect of atmospheric radiance errors in radiometric sea surface
skin temperature measurements. Applied Optics 39: 2392--2397. Jessup AT, Zappa CJ, and Yeh H (1997) Defining and quantifying micro-scale wave breaking with infrared imagery. Journal of Geophysical Research 102: 23145--23153. McKeown W and Asher W (1997) A radiometric method to measure the concentration boundary layer thickness at an air–water interface. Journal of Atmospheric and Oceanic Technology 14: 1494--1501. Shaw JA (1999) Degree of polarisation in spectral radiances from water viewing infrared radiometers. Applied Optics 15: 3157--3165. Smith WL, Knuteson RO, Rivercombe HH, et al. (1996) Observations of the infrared radiative properties of the ocean – implications for the measurement of sea surface temperature via satellite remote sensing. Bulletin of the American Meteorological Society 77: 41--51. Suarez MJ, Emery WJ, and Wick GA (1997) The multichannel infrared sea truth radiometric calibrator (MISTRC). Journal of Atmospheric and Oceanic Technology 14: 243--253. Thomas JP, Knight RJ, Roscoe HK, Turner J, and Symon C (1995) An evaluation of a self-calibrating infrared radiometer for measuring sea surface temperature. Journal of Atmospheric and Oceanic Technology 12: 301--316.
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IRON FERTILIZATION K. H. Coale, Moss Landing Marine Laboratories, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1385–1397, & 2001, Elsevier Ltd.
Introduction
Light 0 Surface mixed layer Euphotic zone 50 Depth (m)
The trace element iron has been shown to play a critical role in nutrient utilization and phytoplankton growth and therefore in the uptake of carbon dioxide from the surface waters of the global ocean. Carbon fixation in the surface waters, via phytoplankton growth, shifts the ocean–atmosphere exchange equilibrium for carbon dioxide. As a result, levels of atmospheric carbon dioxide (a greenhouse gas) and iron flux to the oceans have been linked to climate change (glacial to interglacial transitions). These recent findings have led some to suggest that largescale iron fertilization of the world’s oceans might therefore be a feasible strategy for controlling climate. Others speculate that such a strategy could deleteriously alter the ocean ecosystem, and still others have calculated that such a strategy would be ineffective in removing sufficient carbon dioxide to produce a sizable and rapid result. This article focuses on carbon and the major plant nutrients, nitrate, phosphate, and silicate, and describes how our recent discovery of the role of iron in the oceans has increased our understanding of phytoplankton growth, nutrient cycling, and the flux of carbon from the atmosphere to the deep sea.
ratio (Redfield, 1934, 1958) and can be expressed on a molar basis relative to carbon as 106C : 16N : 1P. Significant local variations in this uptake/regeneration relationship can be found and are a function of the phytoplankton community and growth conditions, yet this ratio can serve as a conceptual model for nutrient uptake and export. The vertical distribution of the major nutrients typically shows surface water depletion and increasing concentrations with depth. The schematic profile in Figure 1 reflects the processes of phytoplankton uptake within the euphotic zone and remineralization of sinking planktonic debris via microbial degradation, leading to increased concentrations in the deep sea. Given favorable growth conditions, the nutrients at the surface may be depleted to zero. The rate of phytoplankton production of new biomass, and therefore the rate of carbon uptake, is controlled by the resupply of nutrients to the surface waters, usually via the upwelling of deep waters. Upwelling occurs over the entire ocean basin at the rate of approximately 4 m per year but increases in coastal and
Major Nutrients Phytoplankton growth in the oceans requires many physical, chemical, and biological factors that are distributed inhomogenously in space and time. Because carbon, primarily in the form of the bicarbonate ion, and sulfur, as sulfate, are abundant throughout the water column, the major plant nutrients in the ocean commonly thought to be critical for phytoplankton growth are those that exist at the micromolar level such as nitrate, phosphate, and silicate. These, together with carbon and sulfur, form the major building blocks for biomass in the sea. As fundamental cellular constituents, they are generally thought to be taken up and remineralized in constant ratio to one another. This is known as the Redfield
100
Temperature NO3
–
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200 Figure 1 A schematic profile indicating the regions of the upper water column where phytoplankton grow. The surface mixed layer is that region that is actively mixed by wind and wave energy, which is typically depleted in major nutrients. Below this mixed layer temperatures decrease and nutrients increase as material sinking from the mixed layer is regenerated by microbial decomposition.
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regions of divergent surface water flow, reaching average values of 15 to 30 or greater. Thus, those regions of high nutrient supply or persistent high nutrient concentrations are thought to be most important in terms of carbon removal.
Nitrogen versus Phosphorus Limitation Although both nitrogen and phosphorus are required at nearly constant ratios characteristic of deep water, nitrogen has generally been thought to be the limiting nutrient in sea water rather than phosphorus. This idea has been based on two observations: selective enrichment experiments and surface water distributions. When ammonia and phosphate are added to sea water in grow-out experiments, phytoplankton growth increases with the ammonia addition and not with the phosphate addition, thus indicating that reduced nitrogen and not phosphorus is limiting. Also, when surface water concentration of nitrate and phosphate are plotted together (Figure 2), it appears that there is still residual phosphate after the nitrate has gone to zero.
The notion of nitrogen limitation seems counterintuitive when one considers the abundant supply of dinitrogen (N2) in the atmosphere. Yet this nitrogen gas is kinetically unavailable to most phytoplankton because of the large amount of energy required to break the triple bond that binds the dinitrogen molecule. Only those organisms capable of nitrogen fixation can take advantage of this form of nitrogen and reduce atmospheric N2 to biologically available nitrogen in the form of urea and ammonia. This is, energetically, a very expensive process requiring specialized enzymes (nitrogenase), an anaerobic microenvironment, and large amounts of reducing power in the form of electrons generated by photosynthesis. Although there is currently the suggestion that nitrogen fixation may have been underestimated as an important geochemical process, the major mode of nitrogen assimilation, giving rise to new plant production in surface waters, is thought to be nitrate uptake. The uptake of nitrate and subsequent conversion to reduced nitrogen in cells requires a change of five in the oxidation state and proceeds in a stepwise fashion. The initial reduction takes place via the nitrate/nitrite reductase enzyme present in phytoplankton and requires large amounts of the reduced nicotinamide– adenine dinucleotide phosphate (NADPH) and of adenosine triphosphate (ATP) and thus of harvested light energy from photosystem II. Both the nitrogenase enzyme and the nitrate reductase enzyme require iron as a cofactor and are thus sensitive to iron availability.
Ocean Regions
Figure 2 A plot of the global surface water concentrations of phosphate versus nitrate indicating a general positive intercept for phosphorus when nitrate has gone to zero. This is one of the imperical observations favoring the notion of nitrate limitation over phosphate limitation.
From a nutrient and biotic perspective, the oceans can be generally divided into biogeochemical provinces that reflect differences in the abundance of macronutrients and the standing stocks of phytoplankton. These are the high-nitrate, high-chlorophyll (HNHC); high-nitrate, low-chlorophyll (HNLC); low-nitrate, high-chlorophyll (LNHC); and low-nitrate, low-chlorophyll (LNLC) regimes (Table 1). Only the HNLC and LNLC regimes are relatively stable, because the high phytoplankton
Table 1 The relationship between biomass and nitrate as a function of biogeochemical province and the approximate ocean area represented by these regimes High-chlorophyll
Low-chlorophyll
High-nitrate
Unstable/coastal (5%)
Low-nitrate
Unstable/coastal (5%)
Stable/Subarctic/Antarctic/ equatorial Pacific (20%) Oligotrophic gyres (70%)
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Figure 3 A schematic representation of the ‘iron theory’ as it functions in offshore HNLC regions and coastal transient LNHC regions. It has been suggested that iron added to the HNLC regions would induce them to function as LNHC regions and promote carbon export.
growth rates in the other two systems will deplete any residual nitrate and sink out of the system. The processes that give rise to these regimes have been the subject of some debate over the last few years and are of fundamental importance relative to carbon export (Figure 3). High-nitrate, Low-chlorophyll Regions
The HNLC regions are thought to represent about 20% of the areal extent of the world’s oceans. These are generally regions characterized by more than 2 mmol l 1 nitrate and less than 0.5 mg l 1 chlorophyll-a, a proxy for plant biomass. The major HNLC regions are shown in Figure 4 and represent the Subarctic Pacific, large regions of the eastern equatorial Pacific and the Southern Ocean. These HNLC regions persist in areas that have high macronutrient concentrations, adequate light, and physical characteristics required for phytoplankton growth but have very low plant biomass. Two explanations have been
given to describe the persistence of this condition. (1) The rates of zooplankton grazing of the phytoplankton community may balance or exceed phytoplankton growth rates in these areas, thus cropping plant biomass to very low levels and recycling reduced nitrogen from the plant community, thereby decreasing the uptake of nitrate. (2) Some other micronutrient (possibly iron) physiologically limits the rate of phytoplankton growth. These are known as top-down and bottom-up control, respectively. Several studies of zooplankton grazing and phytoplankton growth in these HNLC regions, particularly the Subarctic Pacific, confirm the hypothesis that grazers control production in these waters. Recent physiological studies, however, indicate that phytoplankton growth rates in these regions are suboptimal, as is the efficiency with which phytoplankton harvest light energy. These observations indicate that phytoplankton growth may be limited by something other than (or in addition to) grazing. Specifically, these studies implicate the lack of
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Figure 4 Current HNLC regions of the world’s oceans covering an extimated 20% of the ocean surface. These regions include the Subarctic Pacific, equatorial Pacific and Southern Ocean.
sufficient electron transport proteins and the cell’s ability to transfer reducing power from the photocenter. These have been shown to be symptomatic of iron deficiency.
The Role of Iron Iron is a required micronutrient for all living systems. Because of its d-electron configuration, iron readily undergoes redox transitions between Fe(II) and Fe(III) at physiological redox potentials. For this reason, iron is particularly well suited to many enzyme and electron carrier proteins. The genetic sequences coding for many iron-containing electron carriers and enzymes are highly conserved, indicating iron and iron-containing proteins were key features of early biosynthesis. When life evolved, the atmosphere and waters of the planet were reducing and iron was abundant in the form of soluble Fe(II). Readily available and at high concentration, iron was not likely to have been limiting in the primordial biosphere. As photosynthesis evolved, oxygen was produced as a by-product. As the biosphere became more oxidizing, iron precipitated from aquatic systems in vast quantities, leaving phytoplankton and other aquatic life forms in a vastly changed and newly deficient chemical milieu. Evidence of this mass Fe(III) precipitation event is captured in the ancient banded iron formations in many parts of the world. Many primitive aquatic and terrestrial
organisms have subsequently evolved the ability to sequester iron through the elaboration of specific Fe(II)-binding ligands, known as siderophores. Evidence for siderophore production has been found in several marine dinoflagellates and bacteria and some researchers have detected similar compounds in sea water. Today, iron exists in sea water at vanishingly small concentrations. Owing to both inorganic precipitation and biological uptake, typical surface water values are on the order of 20 pmol l 1, perhaps a billion times less than during the prehistoric past. Iron concentrations in the oceans increase with depth, in much the same manner as the major plant nutrients (Figure 5). The discovery that iron concentrations in surface waters is so low and shows a nutrient-like profile led some to speculate that iron availability limits plant growth in the oceans. This notion has been tested in bottle enrichment experiments throughout the major HNLC regions of the world’s oceans. These experiments have demonstrated dramatic phytoplankton growth and nutrient uptake upon the addition of iron relative to control experiments in which no iron was added. Criticism that such small-scale, enclosed experiments may not accurately reflect the response of the HNLC system at the level of the community has led to several large-scale iron fertilization experiments in the equatorial Pacific and Southern Ocean. These have been some of the most dramatic
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IRON FERTILIZATION
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developed in four release experiments in the equatorial Pacific (IronEx I and II) and more recently in the Southern Ocean (SOIREE). At this writing, a similar strategy is being employed in the Caruso experiments now underway in the Atlantic sector of the Southern Ocean. All of these strategies were developed to address certain scientific questions and were not designed as preliminary to any geoengineering effort. Form of Iron
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Figure 5 The vertical distributions of iron, nitrate, silicate, and oxygen in sea water. This figure shows how iron is depleted to picomolar levels in surface waters and has a profile that mimics other plant nutrients.
oceanographic experiments of our times and have led to a profound and new understanding of ocean systems.
Open Ocean Iron Enrichment The question of iron limitation was brought into sharp scientific focus with a series of public lectures, reports by the US National Research Council, papers, special publications, and popular articles between 1988 and 1991. What was resolved was the need to perform an open ocean enrichment experiment in order to definitively test the hypothesis that iron limits phytoplankton growth and nutrient and carbon dioxide uptake in HNLC regions. Such an experiment posed severe logistical challenges and had never been conducted.
Experimental Strategy The mechanics of producing an iron-enriched experimental patch and following it over time was
All experiments to date have involved the injection of an iron sulfate solution into the ship’s wake to achieve rapid dilution and dispersion throughout the mixed layer (Figure 6). The rationale for using ferrous sulfate involved the following considerations: (1) ferrous sulfate is the most likely form of iron to enter the oceans via atmospheric deposition; (2) it is readily soluble (initially); (3) it is available in a relatively pure form so as to reduce the introduction of other potentially bioactive trace metals; and (4) its counterion (sulfate) is ubiquitous in sea water and not likely to produce confounding effects. Although mixing models indicate that Fe(II) carbonate may reach insoluble levels in the ship’s wake, rapid dilution reduces this possibility. New forms of iron are now being considered by those who would seek to reduce the need for subsequent infusions. Such forms could include iron lignosite, which would increase the solubility and residence time of iron in the surface waters. Since this is a chelated form of iron, problems of rapid precipitation are reduced. In addition, iron lignosulfonate is about 15% Fe by weight, making it a space-efficient form of iron to transport. As yet untested is the extent to which such a compound would reduce the need for re-infusion. Although solid forms of iron have been proposed (slow-release iron pellets; finely milled magnetite or iron ores), the ability to trace the enriched area with an inert tracer has required that the form of iron added and the tracer both be in the dissolved form. Inert Tracer
Concurrent with the injection of iron is the injection of the inert chemical tracer sulfur hexafluoride (SF6). By presaturating a tank of sea water with SF6 and employing an expandable displacement bladder, a constant molar injection ratio of Fe : SF6 can be achieved (Figure 6). In this way, both conservative and nonconservative removal of iron can be quantified. Sulfur hexafluoride traces the physical properties of the enriched patch; the relatively rapid shipboard detection of SF6 can be used to track and
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Figure 6 The iron injection system used during the IronEx experiments utilized two polyethylene tanks that could be sequentially filled with sea water and iron sulfate solution while the other was being injected behind the ship’s propellers. A steel tank of sea water saturated with 40 g of sulfur hexafluoride (SF6) was simultaneously mixed with the iron sulfate solution to provide a conservative tracer of mixing.
map the enriched area. The addition of helium-3 to the injected tracer can provide useful information regarding gas transfer. Fluorometry
The biophysical response of the phytoplankton is rapid and readily detectable. Thus shipboard measurement of relative fluorescence (Fv/Fm) using fast repetition rate fluorometry has been shown to be a useful tactical tool and gives nearly instantaneous mapping and tracking feedback.
Remote Sensing
A variety of airborne and satellite-borne active and passive optical packages provide rapid, large-scale mapping and tracking of the enriched area. Although SeaWiffs was not operational during IronEx I and II, AVHRR was able to detect the IronEx II bloom and airborne optical LIDAR was very useful during IronEx I. SOIREE has made very good use of the more recent SeaWiffs images, which have markedly extended the observational period and led to new hypotheses regarding iron cycling in polar systems.
Shipboard Iron Analysis
Experimental Measurements
Because iron is rapidly lost from the system (at least initially), the shipboard determination of iron is necessary to determine the timing and amount of subsequent infusions. Several shipboard methods, using both chemiluminescent and catalytic colorimetric detection have proven useful in this regard.
In addition to the tactical measurements and remote sensing techniques required to track and ascertain the development of the physical dynamics of the enriched patch, a number of measurements have been made to track the biogeochemical development of the experiment. These have typically involved a series of underway measurements made using the ship’s flowing sea water system or towed fish. In addition, discrete measurements are made in the vertical dimension at every station occupied both inside and outside of the fertilized area. These measurements include temperature salinity, fluorescence (a measure of plant biomass), transmissivity (a measure of suspended particles), oxygen, nitrate,
Lagrangian Drifters
A Lagrangian point of reference has proven to be very useful in every experiment to date. Depending upon the advective regime, this is the only practical way to achieve rapid and precise navigation and mapping about the enriched area.
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IRON FERTILIZATION
phosphate, silicate, carbon dioxide partial pressure, pH, alkalinity, total carbon dioxide, iron-binding ligands, 234Th : 238U radioisotopic disequilibria (a proxy for particle removal), relative fluorescence (indicator of photosynthetic competence), primary production, phytoplankton and zooplankton enumeration, grazing rates, nitrate uptake, and particulate and dissolved organic carbon and nitrogen. These parameters allow for the general characterization of both the biological and geochemical response to added iron. From the results of the equatorial enrichment experiments (IronEx I and II) and the Southern Ocean Iron Enrichment Experiment (SOIREE), several general features have been identified.
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unfavorable redox process is only made possible by substantial reducing power (in the form of NADPH) made available through photosynthesis and active nitrate reductase, an iron-requiring enzyme. Without iron, plants cannot take up nitrate efficiently. This provided original evidence implicating iron deficiency as the cause of the HNLC condition. When phytoplankton communities are relieved from iron deficiency, specific rates of nitrate uptake increase. This has been observed in both the equatorial Pacific and the Southern Ocean using isotopic tracers of nitrate uptake and conversion. In addition, the accelerated uptake of nitrate has been observed in both the mesoscale iron enrichment experiments to date, IronEx and SOIREE. Growth Response
Findings to Date Biophysical Response
The experiments to date have focused on the highnitrate, low-chlorophyll (HNLC) areas of the world’s oceans, primarily in the Subarctic, equatorial Pacific and Southern Ocean. In general, when light is abundant many researchers find that HNLC systems are iron-limited. The nature of this limitation is similar between regions but manifests itself at different levels of the trophic structure in some characteristic ways. In general, all members of the HNLC photosynthetic community are physiologically limited by iron availability. This observation is based primarily on the examination of the efficiency of photosystem II, the light-harvesting reaction centers. At ambient levels of iron, light harvesting proceeds at suboptimal rates. This has been attributed to the lack of iron-dependent electron carrier proteins at low iron concentrations. When iron concentrations are increased by subnanomolar amounts, the efficiency of light harvesting rapidly increases to maximum levels. Using fast repetition rate fluorometry and non-heme iron proteins, researchers have described these observations in detail. What is notable about these results is that iron limitation seems to affect the photosynthetic energy conversion efficiency of even the smallest of phytoplankton. This has been a unique finding that stands in contrast to the hypothesis that, because of diffusion, smaller cells are not iron limited but larger cells are. Nitrate Uptake
As discussed above, iron is also required for the reduction (assimilation) of nitrate. In fact, a change of oxidation state of five is required between nitrate and the reduced forms of nitrogen found in amino acids and proteins. Such a large and energetically
When iron is present, phytoplankton growth rates increase dramatically. Experiments over widely differing oceanographic regimes have demonstrated that, when light and temperature are favorable, phytoplankton growth rates in HNLC environments increase to their maximum at dissolved iron concentrations generally below 0.5 nmol l 1. This observation is significant in that it indicates that phytoplankton are adapted to very low levels of iron and they do not grow faster if given iron at more than 0.5 nmol l 1. Given that there is still some disagreement within the scientific community about the validity of some iron measurements, this phytoplankton response provides a natural, environmental, and biogeochemical benchmark against which to compare results. The iron-induced transient imbalance between phytoplankton growth and grazing in the equatorial Pacific during IronEx II resulted in a 30-fold increase in plant biomass (Figure 7). Similarly, a 6-fold increase was observed during the SOIREE experiment in the Southern Ocean. These are perhaps the most dramatic demonstrations of iron limitation of nutrient cycling, and phytoplankton growth to date and has fortified the notion that iron fertilization may be a useful strategy to sequester carbon in the oceans. Heterotrophic Community
As the primary trophic levels increase in biomass, growth in the small microflagellate and heterotrophic bacterial communities increase in kind. It appears that these consumers of recently fixed carbon (both particulate and dissolved) respond to the food source and not necessarily the iron (although some have been found to be iron-limited). Because their division rates are fast, heterotrophic bacteria, ciliates, and flagellates can rapidly divide and
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Figure 7 Chlorophyll concentrations during IronEx II were mapped daily. This figure shows the progression of the phytoplankton bloom that reached over 30 times the background concentrations.
Nutrient Uptake Ratios
3 H
SiO4 / NO3–
respond to increasing food availability to the point where the growth rates of the smaller phytoplankton can be overwhelmed by grazing. Thus there is a much more rapid turnover of fixed carbon and nitrogen in iron replete systems. M. Landry and coworkers have documented this in dilution experiments conducted during IronEx II. These results appear to be consistent with the recent SOIREE experiments as well.
SiO4: NO3 uptake ratio vs. dissolved iron concentration
J H 2 H HJJ J J H J J JJ H JH JH J H H H 1
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An imbalance in production and consumption, however, can arise at the larger trophic levels. Because the reproduction rates of the larger micro- and mesozooplankton are long with respect to diatom division rates, iron-replete diatoms can escape the pressures of grazing on short timescales (weeks). This is thought to be the reason why, in every iron enrichment experiment, diatoms ultimately dominate in biomass. This result is important for a variety of reasons. It suggests that transient additions of iron would be most effective in producing net carbon uptake and it implicates an important role of silicate in carbon flux. The role of iron in silicate uptake has been studied extensively by Franck and colleagues. The results, together with those of Takeda and coworkers, show that iron alters the uptake ratio of nitrate and silicate at very low levels (Figure 8). This is thought to be brought about by the increase in nitrate uptake rates relative to silica.
0
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Figure 8 Bottle enrichment experiments show that the silicate : nitrate uptake ratio changes as a function of the iron added. This is thought to be due to the increased rate of iron uptake relative to silicate in these experimental treatments.
Organic Ligands
Consistent with the role of iron as a limiting nutrient in HNLC systems is the notion that organisms may have evolved competitive mechanisms to increase iron solubility and uptake. In terrestrial systems this is accomplished using extracellularly excreted or membrane-bound siderophores. Similar compounds have been shown to exist in sea water where the competition for iron may be as fierce as it is on land. In open ocean systems where it has been measured, iron-binding ligand production increases with the
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IRON FERTILIZATION
addition of iron. Whether this is a competitive response to added iron or a function of phytoplankton biomass and grazing is not yet well understood. However, this is an important natural mechanism for reducing the inorganic scavenging of iron from the surface waters and increasing iron availability to phytoplankton. More recent studies have considerably advanced our understanding of these ligands, their distribution and their role in ocean ecosystems. Carbon Flux
It is the imbalance in the community structure that gives rise to the geochemical signal. Whereas iron stimulation of the smaller members of the community may result in chemical signatures such as an increased production of beta-dimethylsulfonioproprionate (DMSP), it is the stimulation of the larger producers that decouples the large cell producers from grazing and results in a net uptake and export of nitrate, carbon dioxide, and silicate. The extent to which this imbalance results in carbon flux, however, has yet to be adequately described. The inability to quantify carbon export has primarily been a problem of experimental scale. Even though mesoscale experiments have, for the first time, given us the ability to address the effect of iron on communities, the products of surface water processes and the effects on the midwater column have been difficult to track. For instance, in the IronEx II experiment, a time-series of the enriched patch was diluted by 40% per day. The dilution was primarily in a lateral
339
(horizontal/isopycnal) dimension. Although some correction for lateral dilution can be made, our ability to quantify carbon export is dependent upon the measurement of a signal in waters below the mixed layer or from an uneroded enriched patch. Current data from the equatorial Pacific showed that the IronEx II experiment advected over six patch diameters per day. This means that at no time during the experiment were the products of increased export reflected in the waters below the enriched area. A transect through the IronEx II patch is shown in Figure 9. This figure indicates the massive production of plant biomass with a concomitant decrease in both nitrate and carbon dioxide. The results from the equatorial Pacific, when corrected for dilution, suggest that about 2500 t of carbon were exported from the mixed layer over a 7day period. These results are preliminary and subject to more rigorous estimates of dilution and export production, but they do agree favorably with estimates based upon both carbon and nitrogen budgets. Similarly, thorium export was observed in this experiment, confirming some particle removal. The results of the SOIREE experiment were similar in many ways but were not as definitive with respect to carbon flux. In this experiment biomass increased 6-fold, nitrate was depleted by 2 mmol l 1 and carbon dioxide by 35–40 microatmospheres (3.5–4.0 Pa). This was a greatly attenuated signal relative to IronEx II. Colder water temperatures likely led to slower rates of production and bloom evolution and there was no observable carbon flux.
Figure 9 A transect through the IronEx II patch. The x-axis shows GMT as the ship steams from east to west through the center of the patch. Simultaneously plotted are the iron-induced production of chlorophyll, the drawdown of carbon dioxide, and the uptake of nitrate in this bloom.
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IRON FERTILIZATION
Figure 10 Simple calculations of the potential for carbon export for the Southern Ocean. These calculations are based on the necessary amount of iron required to efficiently utilize the annual upwelled nitrate and the subsequent incorporation into sinking organic matter. An estimated 1.8 109 t (Gt) of carbon export could be realized in this simple model.
Original estimates of carbon export in the Southern Ocean based on the iron-induced efficient utilization of nitrate suggest that as much as 1.8 109 t of carbon could be removed annually (Figure 10). These estimates of carbon sequestration have been challenged by some modelers yet all models lack important experimental parameters which will be measured in upcoming experiments.
Remaining Questions A multitude of questions remain regarding the role of iron in shaping the nature of the pelagic community. The most pressing question is whether iron enrichment accelerates the downward transport of carbon from the surface waters to the deep sea? More specifically, how does iron affect the cycling of carbon in HNLC, LNLC, and coastal systems? Recent studies indicate that coastal systems may be ironlimited and the iron requirement for nitrogenase activity is quite large, suggesting that iron may limit nitrogen fixation, but there have been limited studies to test the former and none to test the latter. If iron does stimulate carbon uptake, what are the spatial scales over which this fixed carbon may be remineralized? This is crucial to predicting whether fertilization is an effective carbon sequestration mechanism. Given these considerations, the most feasible way to understand and quantify carbon export from an enriched water mass is to increase the scale of the experiment such that both lateral dilution and submixed-layer relative advection are small with respect to the size of the enriched patch. For areas such as
the equatorial Pacific, this would be very large (hundreds of kilometers on a side). For other areas, it could be much smaller. The focus of the IronEx and SOIREE experiments has been from the scientific perspective, but this focus is shifting toward the application of iron enrichment as a carbon sequestration strategy. We have come about rapidly from the perspective of trying to understand how the world works to one of trying to make the world work for us. Several basic questions remain regarding the role of natural or anthropogenic iron fertilization on carbon export. Some of the most pressing questions are: What are the best proxies for carbon export? How can carbon export best be verified? What are the long-term ecological consequences of iron enrichment on surface water community structure, midwater processes, and benthic processes? Even with answers to these, there are others that need to be addressed prior to any serious consideration of iron fertilization as an ocean carbon sequestration option. Simple technology is sufficient to produce a massive bloom. The technology required either for a large-scale enrichment experiment or for purposeful attempts to sequester carbon is readily available. Ships, aircraft (tankers and research platforms), tracer technology, a broad range of new Autonomous Underwater Vehicles (AUVs) and instrument packages, Lagrangian buoy tracking systems, together with aircraft and satellite remote sensing systems and a new suite of chemical sensors/in situ detection technologies are all available, or are being developed. Industrial bulk handling equipment is available for large-scale implementation. The big questions, however, are larger than the technology. With a slow start, the notion of both scientific experimentation through manipulative experiments, as well as the use of iron to purposefully sequester carbon, is gaining momentum. There are now national, international, industrial, and scientific concerns willing to support larger-scale experiments. The materials required for such an experiment are inexpensive and readily available, even as industrial by-products (of paper, mining, and steel processing). Given the concern over climate change and the rapid modernization of large developing countries such as China and India, there is a pressing need to address the increased emission of greenhouse gases. Through the implementation of the Kyoto accords or other international agreements to curb emissions (Rio), financial incentives will reach into the multibillion dollar level annually. Certainly there will soon be an overwhelming fiscal incentive to investigate, if not implement, purposeful open ocean carbon sequestration trials.
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IRON FERTILIZATION
341
A Societal Challenge
Further Reading
The question is not whether we have the capability of embarking upon such an engineering strategy but whether we have the collective wisdom to responsibly negotiate such a course of action. Posing the question another way: If we do not have the social, political and economic tools or motivation to control our own population and greenhouse gas emissions, what gives us the confidence that we have the wisdom and ability to responsibly manipulate and control large ocean ecosystems without propagating yet another massive environmental calamity? Have we as an international community first tackled the difficult but obvious problem of overpopulation and implemented alternative energy technologies for transportation, industry, and domestic use? Other social questions arise as well. Is it appropriate to use the ocean commons for such a purpose? What individuals, companies, or countries would derive monetary compensation for such an effort and how would this be decided? It is clear that there are major scientific investigations and findings that can only benefit from largescale open ocean enrichment experiments, but certainly a large-scale carbon sequestration effort should not proceed without a clear understanding of both the science and the answers to the questions above.
Abraham ER, Law CS, Boyd PW, et al. (2000) Importance of stirring in the development of an iron-fertilized phytoplankton bloom. Nature 407: 727--730. Barbeau K, Moffett JW, Caron DA, Croot PL, and Erdner DL (1996) Role of protozoan grazing in relieving iron limitation of phytoplankton. Nature 380: 61--64. Behrenfeld MJ, Bale AJ, Kobler ZS, Aiken J, and Falkowski PG (1996) Confirmation of iron limitation of phytoplankton photosynthesis in Equatorial Pacific Ocean. Nature 383: 508--511. Boyd PW, Watson AJ, Law CS, et al. (2000) A mesoscale phytoplankton bloom in the polar Southern Ocean stimulated by iron fertilization. Nature 407: 695-702. Cavender-Bares KK, Mann EL, Chishom SW, Ondrusek ME, and Bidigare RR (1999) Differential response of equatorial phytoplankton to iron fertilization. Limnology and Oceanography 44: 237--246. Coale KH, Johnson KS, Fitzwater SE, et al. (1996) A massive phytoplankton bloom induced by an ecosystemscale iron fertilization experiment in the equatorial Pacific Ocean. Nature 383: 495--501. Coale KH, Johnson KS, Fitzwater SE, et al. (1998) IronExI, an in situ iron-enrichment experiment: experimental design, implementation and results. Deep-Sea Research Part II 45: 919--945. Elrod VA, Johnson KS, and Coale KH (1991) Determination of subnanomolar levels of iron (II) and total dissolved iron in seawater by flow injection analysis with chemiluminescence dection. Analytical Chemistry 63: 893--898. Fitzwater SE, Coale KH, Gordon RM, Johnson KS, and Ondrusek ME (1996) Iron deficiency and phytoplankton growth in the equatorial Pacific. Deep-Sea Research Part II 43: 995--1015. Greene RM, Geider RJ, and Falkowski PG (1991) Effect of iron lititation on photosynthesis in a marine diatom. Limnology Oceanogrography 36: 1772--1782. Hoge EF, Wright CW, Swift RN, et al. (1998) Fluorescence signatures of an iron-enriched phytoplankton community in the eastern equatorial Pacific Ocean. Deep-Sea Research Part II 45: 1073--1082. Johnson KS, Coale KH, Elrod VA, and Tinsdale NW (1994) Iron photochemistry in seawater from the Equatorial Pacific. Marine Chemistry 46: 319--334. Kolber ZS, Barber RT, Coale KH, et al. (1994) Iron limitation of phytoplankton photosynthesis in the Equatorial Pacific Ocean. Nature 371: 145--149. Landry MR, Ondrusek ME, Tanner SJ, et al. (2000) Biological response to iron fertilization in the eastern equtorial Pacific (Ironex II). I. Microplankton community abundances and biomass. Marine Ecology Progress Series 201: 27--42. LaRoche J, Boyd PW, McKay RML, and Geider RJ (1996) Flavodoxin as an in situ marker for iron stress in phytoplankton. Nature 382: 802--805. Law CS, Watson AJ, Liddicoat MI, and Stanton T (1998) Sulfer hexafloride as a tracer of biogeochemical and
Glossary ATP AVHRR
Adenosine triphosphate Advanced Very High Resolution Radiometer HNHC High-nitrate high-chlorophyll HNLC High-nitrate low-chlorophyll IronEx Iron Enrichment Experiment LIDAR Light detection and ranging LNHC Low-nitrate high-chlorophyll LNLC Low-nitrate low-chlorophyll NADPH Reduced form of nicotinamide–adenine dinucleotide phosphate SOIREE Southern Ocean Iron Enrichment Experiment
See also Fluorometry for Biological Sensing. Fluorometry for Chemical Sensing. Nitrogen Cycle. Phosphorus Cycle. Platforms: Autonomous Underwater Vehicles. Primary Production Distribution. Primary Production Processes. Redfield Ratio. Satellite Remote Sensing SAR.
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physical processes in an open-ocean iron fertilization experiment. Deep-Sea Research Part II 45: 977--994. Martin JH, Coale KH, Johnson KS, et al. (1994) Testing the iron hypothesis in ecosystems of the equatorial Pacific Ocean. Nature 371: 123--129. Nightingale PD, Liss PS, and Schlosser P (2000) Measurements of air–gas transfer during an open ocean algal bloom. Geophysical Research Letters 27: 2117--2121. Obata H, Karatani H, and Nakayama E (1993) Automated determination of iron in seawater by chelating resin concentration and chemiluminescence detection. Analytical Chemistry 65: 1524--1528. Redfield AC (1934) On the proportions of organic derivatives in sea water and their relation to the composition of plankton. James Johnstone Memorial Volume, pp. 177--192. Liverpool: Liverpool University Press. Redfield AC (1958) The biological control of chemical factors in the environment. American Journal of Science 46: 205--221. Rue EL and Bruland KW (1997) The role of organic complexation on ambient iron chemistry in the equatorial Pacific Ocean and the response of a mesoscale iron addition experiment. Limnology and Oceanography 42: 901--910. Smith SV (1984) Phosphorus versus nitrogen limitation in the marine environment. Limnology and Oceanography 29: 1149--1160.
Stanton TP, Law CS, and Watson AJ (1998) Physical evolutation of the IronEx I open ocean tracer patch. Deep-Sea Research Part II 45: 947--975. Takeda S and Obata H (1995) Response of equatorial phytoplankton to subnanomolar Fe enrichment. Marine Chemistry 50: 219--227. Trick CG and Wilhelm SW (1995) Physiological changes in coastal marine cyanobacterium Synechococcus sp. PCC 7002 exposed to low ferric ion levels. Marine Chemistry 50: 207--217. Turner SM, Nightingale PD, Spokes LJ, Liddicoat MI, and Liss PS (1996) Increased dimethyl sulfide concentrations in seawater from in situ iron enrichment. Nature 383: 513--517. Upstill-Goddard RC, Watson AJ, Wood J, and Liddicoat MI (1991) Sulfur hexafloride and helium-3 as sea-water tracers: deployment techniques and continuous underway analysis for sulphur hexafloride. Analytica Chimica Acta 249: 555--562. Van den Berg CMG (1995) Evidence for organic complesation of iron in seawater. Marine Chemistry 50: 139--157. Watson AJ, Liss PS, and Duce R (1991) Design of a smallscale in situ iron fertilization experiment. Limnology and Oceanography 36: 1960--1965.
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ISLAND WAKES E. D. Barton, University of Wales, Bangor, Menai Bridge, Anglesey, UK
disturbance or wake is related to the value of the Reynolds number
Copyright & 2001 Elsevier Ltd.
Re ¼
This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1397–1403, & 2001, Elsevier Ltd.
Introduction The ‘island mass effect’ has been documented for about half a century. This refers to a biological enrichment around oceanic islands in comparison to surrounding waters. Despite a relatively large body of evidence to support the existence of such an effect in the vicinity of islands, there have been few studies to investigate the underlying physical causes of this phenomenon. From a physical point of view the presence of an island in the background flow will disturb the flow regime to produce perturbations that ultimately must have biological consequences. Two types of island disturbance have been investigated. The first takes place in shallow, stratified shelf seas with significant tidal regimes but no appreciable mean flow, where as the tide moves water back and forth the island acts as a stirring rod to enhance vertical mixing locally and break down the pycnocline. The second occurs in both shallow and deep water, where flow past an island generates eddies downstream and a wake of disturbed flow extends several island diameters away. This arises in the case of larger islands when a clear ambient flow dominates over tidal variability and also around islands small enough that the tidal stream itself can generate a similar effect. The scale of the eddies is typically close to the island diameter and their time scale will be several days for larger islands but only hours in the case of tidal flows. The nature of the wake and eddies may differ between the oceanic case where conditions can be considered quasigeostrophic and the shallow case where they are frictionally dominated by bottom stress.
Theory Nonrotating Case
The simplest case is where flow past isolated oceanic islands is considered to be analogous to that of channel flow past a circular cylinder. The work of Batchelor presented the case of nonrotating flow in a homogeneous fluid. The form of the downstream
Ud v
½1
where U is the free velocity upstream, d is the diameter of the cylinder and v is the molecular viscosity. Although this formula appears simple, there are certain practical difficulties in applying it to even a nonrotating ocean of homogeneous character. Few islands are isolated or cylindrical, the upstream velocity is generally not well known and finally the molecular viscosity must be replaced by the horizontal eddy viscosity in the ocean, which is in general poorly known. Studies of Aldabra, an Indian Ocean atoll, indicate that the same current speed impinging on the island from different directions produces a different wake because of the asymmetrical form of the island. It is frequently the case that the upstream current in the ocean varies on a range of time scales, or may be subject to horizontal shear, both of which complicate the choice of a suitable value for the free stream velocity. Values of horizontal eddy viscosity coefficients in the ocean based on many experimental determinations increase with the length scale of interest, l, in a nonlinear fashion Kh ðm2 s1 Þ ¼ 2:2 104 l1:13 . Typical values appropriate to oceanic islands vary between 102 and 105 m2 s1. The nature of the wake downstream of the obstacle changes as the Reynolds number increases (Figure 1). For low Reynolds numbers (Reo1) the flow pattern downstream is the same as that upstream and there is no perceptible wake. The flow remains attached to the sides of the cylinder and is laminar throughout the flow field. At Reynolds numbers between 1 and 40, the wake remains basically laminar away from the cylinder and two eddies are formed immediately behind the obstacle where they remain attached. At higher Reynolds number the wake becomes increasingly unstable and counter-rotating eddies form a vortex street. The eddies expand as they move away from the obstacle and gradually decay. For Reynolds numbers Re4 80 eddies formed behind the island no longer remain trapped but are shed alternately into the vortex street. The frequency of eddy shedding n is related to another nondimensional number, the Strouhal number
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St ¼
nd U
½2
343
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ISLAND WAKES
rotation rate of the earth and f the latitude. This variation of the Coriolis parameter f can be represented in terms of the ‘beta plane’ where b ¼ df/dy and y is distance north of the equator. If the dimensionless parameter
Laminar flow
U Re < 1 Attached eddies
Laminar wake
U
b0 ¼
1 < Re < 40
½5
is small, the beta effect may be ignored and the rotation rate taken to be constant on the scale of the island in question. If on the other hand it is of order unity, differences occur from the case of uniform rotation. In particular flow separation is enhanced for flow towards the west and inhibited for flow towards the east.
Attached eddies Unstable wake with eddies
U
bd2 4U
40 < Re < 80
Effects of Bottom Friction
U
Studies of flow patterns around small islands only a few kilometers in diameter in shallow shelf seas indicate that the Reynolds number as defined in eqn[1] overestimates the value at transition between the different cases of wake formation. It has been found that the island wake parameter
Detaching eddies
Re > 80
Figure 1 Reynold number regimes for flow past a cylinder.
This number approaches an asymptotic value of 0.21 at higher Reynolds numbers. Effect of Rotation
The ocean of course is on the rotating Earth and so laboratory experiments have been carried out in rotating tanks to investigate the effect of this rotation. Two dimensionless quantities of importance here are the Rossby number R0 ¼
U Od
½3
which represents the importance of rotation at rate O and the Ekman number 2v Ek ¼ Od2
½4
which determines the width of the wake. The ratio of Rossby number to Ekman number is proportional to the Reynolds number, generalizing this concept to the case of rotating flow. The Earth’s rotation enhances the shedding of eddies in the same sense (cyclonic) so that in the Northern Hemisphere predominantly anticlockwise eddies are to be expected whereas in the Southern Hemisphere clockwise should be more common. Because of the spherical shape of the Earth’s surface the rate of rotation about the local vertical varies with latitude f ¼ 2o sin f, where o is the
P¼
Uh2 Kz d
½6
where U is the stream velocity, h is the water depth, Kz is the vertical eddy diffusivity and d is the dimension of the island, provided better agreement than the Reynolds number between observed and predicted wake parameters. However, P is actually a correct formulation for the Reynolds number when the effect of lateral and bottom frictional boundary layers is taken into account.
Observations Until recently observations of island flow effects had been limited mainly to remote sensing reports of eddy production. Attempts to observe vortex production and development in situ behind oceanic islands have been largely unsuccessful because the background flow has been too weak or variable or the methods of observation insufficient to determine the flow regime adequately. However, a classic early report in 1972, based upon sparse observations of the drift of fishing gear and rudimentary surface current measurement, showed drift patterns downstream of Johnston Atoll in the Pacific Ocean in good agreement with a vortex street situation. In the case of Aldabra Atoll, situated in the South Equatorial Current of the Indian Ocean, surveys made with acoustic Doppler current profiler
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ISLAND WAKES
indicated a single cyclonic (anticlockwise) eddy trapped behind the island in a low Reynolds number regime on two occasions. There was no evidence of continuous eddy or wake production downstream. The flow was variable during the experiment. The first trapped eddy was observed during westward background flow around 20 cm s1. Two subsequent rapid surveys, about ten days apart, of the current field showed predominantly northward and westward flows, respectively. In the first case the free stream flow was about 20 cm s1 and impinged on the island’s widest cross-section. No eddy was observed, only an asymmetrical deviation of currents behind the island. The later case found slightly stronger (30 cm s1) free stream velocity from the east impinging the narrow aspect of the island. This time a second weak eddy of diameter similar to the island width was indicated. Another island showing highly variably flow regimes is Barbados. In Spring 1991, the flow seemed topographically steered around the Barbados ridge and there was no clear evidence of eddy production. The following spring, anticyclonic and cyclonic eddies of similar size to the island were found on either flank (Figure 2). Computer simulations of the flow
14.00
345
regime indicated that typical conditions were conducive to shedding of cyclones and anticyclones alternately with a period of around 10 days. It was not possible to demonstrate the degree to which the observations were attributable to this type of vortex generation, however, and it was concluded that much more extensive observations would be needed to do so. One interesting indication of the simulation was that there was a continuous region of downstream reverse flow towards the island that could provide a return path for fish larvae swept away from the island. Considerable evidence of recurrent eddy shedding has been reported recently in the Canary Island archipelago. There, the Canary Current flows southwestward at an average speed of 5 cm s1. Eddies of both signs have been reported (Figure 3) as being frequently spun off from the island of Gran Canaria. Cyclonic eddies of the same diameter as the island (50 km) and rotation period around 3 days have been observed to develop on the southwestern flank of the island and to move southwest at speeds of 5–15 cm s1. Almost as frequently, anticyclonic eddies have been observed to develop on the south east of the island. They have diameters up to twice
7
20 30
13.75
67
40
65
13.50
61
Latitude
13.25
80 70 60 50 40 30
20
30 L
58
H
50 60
BARBADOS
40
13.00
84
12.75 _ 60.25
_ 60.00
_ 59.75
_ 59.50 Longitude
_ 59.25
_ 59.00
_ 58.75
Figure 2 Sea surface topography in centimeters relative to 250 dbar 25 April–2 May 1991. Arrows denote the direction of surface flow. Note the overall northwestward flow and anticyclonic (H) and cyclonic (L) eddies either side of the island. (Adapted from Bowman et al., 1996.)
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ISLAND WAKES
29N
1607 23 Aug 1999 C
Lee Gran Lee
28N
Canaria
Lee A
Lee C
Lee
Lee
C
27N
A
19W
18W
17W
16W
15W
14W
Figure 3 Sea surface temperature image showing multiple eddies shed from the Canary Islands during August 1999. Cyclones and anticyclones are labeled C and A, respectively.
the size of the cyclonic eddies, similar rotation rates and appear to persist for many months in the region. One such eddy, seeded with drifters in June 1998, was observed to persist for at least seven months (Figure 4). It initially drifted slowly southwestward but later returned northwestward towards the outer islands of the archipelago. Other large anticyclones have been observed downstream of the island pair, Tenerife and La Gomera, trapped close to the islands for at least several weeks before moving away from the island. Smaller cyclones and anticyclones are
29N Tenerife
Fuerteventura Gran Canaria
28N Release point
27N
26N
Jul _ Oct 1998
18W
17W
16W
15W
14W
Figure 4 Path of a drifter released into an anticyclonic eddy shed from Gran Canaria in June 1998. The eddy persisted for 7 months though only the first three months are shown here. (Courtesy of Dr Pablo Sangra, University of Las Palmas de Gran Canaria.)
frequently seen in sea surface temperature images being spun off from the flanks of the other smaller islands. The generation of these eddies may not be entirely a result of the oceanic flow past the islands. The Canaries are high volcanic islands situated in a regime of strong southwestward trade winds. The high island peaks block the flow of the trade winds to form extended lee regions downwind. These are bounded by localized horizontal shear in the wind field and so are locations of strong Ekman divergence and convergence. Ekman transport takes place in the near surface layer and is to the right of the wind in the Northern Hemisphere. At the western boundary of the lee, upwelling of deeper waters must compensate the divergence, while at the eastern boundary, sinking must occur. The upwelling and downwelling elevates or depresses, respectively, the pycnocline from its unperturbed depth. Because the downwind scale of the lee is limited, the elevation or depression of the pycnocline tends to form an eddy of the same sign as expected from the current past the island. The vertical motions expected from the horizontal wind shear on the lee boundaries are of the order of tens of meters per day, and so potentially could contribute significantly to the observed eddy production. Despite the frequent reports of eddies spun off from the islands, it has proven difficult to obtain time series observations of eddy generation which would allow determination of basic eddy properties such as average shedding frequency, size, propagation speed or whether eddies are shed alternately from opposite island flanks. Remote sensing even in the subtropics is often blocked by cloud cover and time series in situ sampling is difficult to maintain.
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ISLAND WAKES
A situation similar to the Canaries has been observed in the case of the Hawaiian archipelago, situated in the North Equatorial Current and trade winds of the Pacific. Downstream of these mountainous islands, the trade winds with speeds of 10– 20 m s1 are separated from the calmer lee by strong boundaries of high wind shear. Locally, the depth of the surface mixed layer depends on wind speed: in the channels between islands, deep mixed layers are observed; in the lee, stirring by the wind is too weak to distribute solar heating down below the surface layer and intense surface warming results during the day. Sharp surface temperature fronts up to 41C, are often associated with these wind shear lines, as is also observed in the Canaries. The lee of islands in both archipelagos is often visible in sea surface temperature images as a significantly warmer triangular area extending downwind. Ekman transports associated with the wind pattern produce pycnocline perturbations as in the Canaries, resulting in intense anticlockwise eddies under the northern shear lines, and less intense clockwise eddies under southern shear lines. The depth of the mixed layer in the lee of Hawaii can vary from less than 20 m in the counterclockwise eddy to more than 120 m in the clockwise eddy. Figure 5 shows a diagram of the Hawaiian island situation which applies equally well to the Canaries. Though the wind has long been viewed as an important generating mechanism for the Hawaiian eddies, it is still unclear how the variability of the wind field affects oceanic eddy generation. The wind itself has often been observed in satellite images of low level clouds to form wakes of counter-rotating atmospheric eddies behind Hawaii and the Canaries. The eddy-shedding period in this case is much
347
shorter, about 10 h, than for oceanic eddies, which have a periodicity of many days. Presumably it is the wind field averaged over the timescale of the ocean eddies that is important. However, it is quite possible that atmospheric eddy shedding is intermittent; during periods when a trapped eddy regime dominates the wind shear lines are relatively stationary and able to feed energy into oceanic eddy production. Further observations are required to unravel the mechanisms at work in these island situations. The North Equatorial Current impinging on the Hawaiian islands, of course, will also tend to produce eddies. The cumulative effect of the many eddies that are spun off is the formation of a mean large scale re-circulation behind the Hawaiian (cf. Barbados) chain that has been named the Hawaii Lee Counter Current. The longevity of individual eddies is further illustrated by one of their surface layer drifters which remained trapped in an anticyclonic vortex formed near the southern flank of the main island, drifting westward over 2000 km at 11 cm s1. This and other drifter tracks showed the remarkable phenomenon of vortex doubling, the process of merging of two vortices of the same sign. When two identicalpffiffiffivortices merge, the radius of the merged eddy is 2 of the original and its period of rotation is doubled. The drifter appeared to show three such vortex-doubling episodes during its trajectory. In the case of larger shallow sea islands in a tidal regime, the flow reverses before a wake can be properly set up. However, the relative motion between the island and surrounding body of water allows the island to act as a ‘stirring rod’ because of the increased flow speeds on the island flanks which produce vertical mixing within a tidal mixing front some distance off the shore. In a stratified region this
Figure 5 Hawaii schematic showing the mechanism of eddy generation by wind shear. Ekman surface layer transports lead to depression and elevation of the pycnocline in regions of convergence and divergence, respectively. This in turn generates oceanic anticyclones and cyclones. (Courtesy of Professor P. Flament, University of Hawaii.)
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ISLAND WAKES
introduces nutrients from the lower layers into the surface layers so enhancing productivity around the island. In the case of the Scilly Islands mixing between low density surface layer and high density bottom layer waters produced intermediate density enriched water spreading out between the layers causing enhancement of the mid-depth chlorophyll maximum in the area around. Similar results were found around St Kilda off western Scotland.
Conclusions Island wakes produced by eddy shedding have been observed in both deep ocean and shallow shelf sea situations. In some cases of shelf sea islands there is no wake as such, because of weak mean flows, but vigorous tidal currents can produce regions of wellmixed water in the surrounding area. In the oceanic case, though there are many examples of eddy and wake observations, there has yet to be a definitive study demonstrating the phenomenon over a range of flow conditions. The effect of islands on the flow regime does have biological consequences, in that enhanced mixing related to shallow sea islands can increase primary production. It has often been argued that island shed eddies may provide a mechanism for retaining fish larvae in the vicinity of their spawning areas. Although the evidence for this is not yet convincing the existence of mean return circulation has been demonstrated in both model and drifter studies of oceanic islands. It was suggested that energy dissipation caused by island flow disturbance could account for 10% of the wind kinetic energy input to the Pacific. However, it is also considered that general turbulent processes known within the deep ocean may be too weak by more than an order of magnitude to explain global redistribution of energy input. In this case vertical mixing must be greater at the ocean boundaries with land than previously considered and the role of islands and island chains may be greater than is presently perceived. Simulations of the Barbados wake indicated that the flow disturbance was extensive, reaching at least eight island diameters downstream.
Both the Canaries and Hawaii archipelagoes clearly control physical conditions for large distances downstream, modifying water masses through mixing, enhancing productivity and shedding long-lived eddies.
See also Canary and Portugal Currents. Ekman Transport and Pumping. Fish Migration, Horizontal. Mesoscale Eddies. Pacific Ocean Equatorial Currents. Upper Ocean Vertical Structure. Wind Driven Circulation.
Further Reading Ari´stegui J, Tett P, Herna´ndez-Guerra A, et al. (1997) The influence of island-generated eddies on chlorophyll distribution: a study of mesoscale variation around Gran Canaria. Deep-Sea Research 44(1): 71–96 Barkley RA (1972) Johnston Atoll’s wake. Journal of Marine Research 30: 201--216. Batchelor GK (1967) An Introduction to Fluid Dynamics. Cambridge: Cambridge University Press. Bowman MJ, Dietrich DE, and Lin CA (1996) Observations and modelling of mesoscale ocean circulation near a small island. In: Maul G (ed.) Small IslandsM: arine Science and Sustainable Development. Coastal and Estuarine Studies, vol. 51, pp. 18--35. Washington: American Geophysical Union. Chopra KP (1973) Atmospheric and oceanic flow problems introduced by islands. Advances in Geophysics 16: 297--421. Flament P, Kennan SC, Lumpkin C, and Stroup ED (1998) The Atlas of Hawai’i. In: Juvik S and Juvik J (eds.) The Ocean, pp. 82--86. Honolulu: University of Hawai’i Press. Simpson JH and Tett P (1986) Island stirring effects on phytoplankton growth. In: Bowman MJ, Yentsch M, and Peterson WT (eds.) Lecture Notes on Coastal and Estuarine Studies, vol. 17, Tidal Mixing and Plankton Dynamics, pp. 41--76. Berlin, Heidelberg: Springer-Verlag. Wolanski E, Imberger J, and Heron ML (1984) Island wakes in shallow coastal waters. Journal of Geophysical Research 89C: 10533--10569.
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KRILL E. J. Murphy, British Antarctic Survey, Marine Life Sciences Division, Cambridge, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1405–1413, & 2001, Elsevier Ltd.
Introduction Krill play a major role in the transfer of energy in marine food webs, being important consumers of phytoplankton and other zooplankton, and prey of many higher trophic level predators that are often commercially important. The importance of krill in the diet of marine predators is reflected in their name; ‘krill’ comes from the Norwegian whaler’s description of the larger food of the great whales. Krill form an order within the Crustacea, the Euphausiacea, which comprises over 80 species in 10 genera. Detailed keys are available to identify individual species that have broadly similar body pattern (Figure 1). The euphausiids occur in a wide range of habitats – coastal, oceanic, and deepoceanregions – and their distributions also extend into the ice-covered regions of the Arctic and the Antarctic. Krill are generally more abundant in higher latitudes and can occur in such large numbers near the surface that they discolor the water. The phylogenetic relationships of many of the euphausiids are unknown, but for some of the key oceanic species their evolutionary development appears to have been associated with the formation of the major circulation patterns of the world’s oceans. This link to large-scale ocean circulation patterns is also reflected in the population distributions and life histories of the euphausiids. Many of the oceanic krill species occur over broad regions in which the centers of the populations tend to be associated with restricted features of the ocean circulation. However, the patterns of flow often result in transport of krill out of their main breeding regions to areas where they do not breed successfully. This also appears to be crucial to their role in many food webs, providing energy input into regions remote from their own main areas of production. The observation that krill are often transported into regions where they do not reproduce also highlights the colonization potential of the group should any changes occur in patterns of ocean circulation. There are several features that mark the euphausiids as unusual plankton. A number of species are
relatively large with a long life span compared with other zooplankton. The largest of the krill grow to over 60 mm and can live for more than 5 years. Another key feature is that in a number of the species the individuals form dense aggregations known asswarms. In some of the larger euphausiids these swarms might more appropriately be thought of as schools, similar to those formed by small fish, where members of the aggregation are aligned and show coherent patterns of behavior. In the Antarctic the term ‘krill’ is often used to denote a single species: the Antarctic krill, Euphausia superba Dana (Figure 1A). This is, as its name suggests, the ‘superb’ krill that is large in size, occurs in vast numbers in the Southern Ocean, and is central to the Antarctic food web. It is the food of not only the now greatly depleted populations of whales but also many of the seals, penguins and other sea birds, and of fish and squid. It is the most studied species and much of the available information on euphausiids in general is based on knowledge of the Antarctic krill, so it is important to remember that this is something of an extreme representative of the group. A number of the euphausiids have been exploited in fisheries. As krill are typically a low trophic level species there has been recognition of the potential impact this could have on the higher trophic levels of marine food webs. The pivotal role of krill in marine food webs has meant that, particularly in the Antarctic, an ecosystem approach to the management of krill fisheries is being developed that has relevance to the sustainable management of marine ecosystems globally.
Species Separation and Geographical Distributions Euphausiids are found throughout the oceans of the world, but their distributions highlight marked differences in habitat and life history amongst apparently similar species. There is a continuing debate about the exact number of species of euphausiids and the degree of separation of subgroups. There are indications from evolutionary studies of mitochondrial DNA that vicariant speciation (separation by formation of a natural barrier) has been important in the development of euphausiid species in the Antarctic. The generation of the Antarctic Polar Front about 25–22 Ma probably led to the separation of
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Figure 1 Two krill species: (A) the Antarctic krill, Euphausia superba and (B) a North Atlantic krill, Meganyctiphanes norvegica.
the ‘Antarctic clade’ (E. superba and E. crystallorophias) from the sister clade of E. vallentini and E. frigida dated at about 20 Ma. Although some euphausiid species occur in coastal and bathypelagic regions (1000–2500 m), most are foundin oceanic epipelagic (0–200 m) and mesopelagic regions (200–1000 m). Although broad distributions have been described for many of the euphausiids, because these animals frequently occur in only relatively low numbers their local distributions are often not well defined. Generally, there is a trend of increasing abundance of krill at higher latitudes. However, there are variations in this pattern, with strong links between the ocean current systems and the regional distribution of krill species. A feature of the euphausiids is that in Southern Ocean and Southern Hemisphere regions many of the key species occur across the full longitudinal range (Figure 2). In the Southern Ocean the ocean circulation is circumpolar, so the same basic pattern of species distribution is found throughout the connected ocean. The key species in the mainly icecovered regions is E. crystallorophias which inhabits the Antarctic continental shelf, although on occasion it has been found transported northward by the major current flows. Further to the north in the seasonally ice-covered areas of the main flow regions of the Antarctic Circumpolar Current are E. superba and E. frigida, with Thysanossa vicina and T. macrura extending northwards to the Antarctic Polar Front. All of these species have heterogeneous distributions in the region. For E. superba there appear to be centers of population in which they can spawn and reproduce successfully, separated by and possibly connected through, regions that are not favorable to breeding but in which krill are found (Figure 3). E. triacantha overlaps the northern limit of E. superba in the south and in the north it overlaps the southern limit of the range of E. vallentini extending north to south of 401S. Further north still are less abundant species suchas E. longirostris and E. lucens that extend north of 401S in areas encompassed by the eastward flows in the southern regions of the main ocean gyres of the Pacific, Atlantic, and Indian Oceans. To the north of this E. similis occurs in all three ocean basin regions, extending from about 50–601S to 301S, but the species is also present further north inthe north-west region of the Indian Ocean to the north of Madagas car. Across the subtropical and tropical regions there is a wide range of species. One that occurs in all the ocean basins is E. tenera, where it has awide distribution but is not very abundant in any region. There are a number of other species found in both the Atlantic and Pacific Oceans. Some species,
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Atlantic only
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M. norvegica T. longicaudata
E. similis E. longirostris E. lucens
Atlantic & Pacific
T. inermis
E. vallentini
T. raschii
E. tenera
E. triachantha
E. gibba group E. superba T. macrura, T. vicina E. frigida E. crystallorophias Pacific only
T. longipes E. pacifica
80˚ N
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40˚ N
0
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40˚ S
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Antarctic
Arctic Figure 2 The broad geographical distributions of some key euphausiid species.
particularly in the central North Pacific and North Atlantic, are abundant but only found in one of the ocean basins. Key species that show this pattern are E. pacifica and T. longipes/T. inspinata that occur only in the northern North Pacific, while in the North Atlantic Meganyctiphanes norvegica and T. longicaudata are dominant (Figure 2). In the northern North Atlantic and North Pacific there are species that occur in both oceans and through into the Arctic regions in the far north. In particular there are two important species, T. raschii and T. inermis, with distributions extending from about 451N to about 801N, although breeding is largely restricted to areas south of 701N. As well as geographical differences there are also marked differences in vertical distribution and many of the species show some form of vertical migration. For example, E. pacifica occurs mainly above 300 m during the day, moving nearer the surface (o150m) at night, while M. norvegica occurs between 100 and 500 m during the day and vertically migrates to shallower depths at night, and in the south
E. superba occurs mainly above about 250 m and migrates nearer the surface at night.
Growth, Development, Physiology Krill species show a range of development strategies that vary between species, and also with the environmental conditions to which they are exposed. Studies of krill population dynamics and development are made difficult because of problems in determining the age of a number of the species. Traditional techniques involving the analysis of the population age structure are still relied upon and a range of mathematical techniques have been employed to distinguish different cohorts in length– frequency size distributions. These are not always definitive and a range of other techniques has been explored such as using age pigment analyses, multiple-morphometric analyses, analyses of structures in the eye, and laboratory maintenance of live specimens. None of these techniques has so far
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120°
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°
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Figure 3 The main regions of occurrence of Antarctic krill E. superba in the Southern Ocean and the pattern of surface circulation fromthe FRAM model (FRAM Group).
provided a good and practical solution to determining the age of krill. However, there is general agreement about the broad characteristics of growthand development of many of the key species. In the Southern Ocean, early studies of E. superba indicated a 2–3 year life cycle based mainly on samplesfrom open-ocean regions. However, further detailed analyses of the size-structure and development of E. superba populations have led to a revision in the life-span up to 45 years with suggestions that in some areas 5–7 year classes can be identified. Laboratory experiments have maintained krill obtained from the sea for 46 years, indicating that a total age of 7–8 years is probably possible in the wild. The development and growth of the krill will depend on the conditions to which they are exposed. E. superba can reach a size of 460 mm with
indications that growth may be very plastic, as in other euphausiid species, varying with the environmental conditions. Thus, krill in more northern and warmer regions may grow more rapidly and develop earlier than krill further south. In these more northern regions, such as around the Island of South Georgia which lies at about 541S, near the Antarctic Polar Front, the krill do not reproduce successfully, with few indications of any viable larvae being found in thearea. The E. superba population in these regions is probably maintained by advection inputs from further south in the Southern Scotia Sea, Weddell Searegion, and from around the Antarctic Peninsula. Such a plastic range of development is illustrated clearly in the northern species T. inermis and T. raschii. These species have a maximum age of
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1 year at the south of their distribution (about 451N), where as further north they survive to over 2 years old, spawning in each year, although females may not mature until their second year. Continuing northward, the maximum age increases and maturation is delayed further with spawning delayed to year 3, and a maximum age of 3 years. In the high Arctic waters the krill still mature but do not spawn and it is the water circulation bringing krill from further south that maintains the species in these areas. Across this range T. inermis grows to over 20 mm, but the rates involved vary depending on the conditions, with slower growth and development occurring further north in their range. M. norvegica is one of the most abundant North Atlantic krill and individuals can reach a maximum size of over 45 mm in some regions. This species shows less age variation across its range than the more northern Thysanossa species, but the variation is still significant. In the south of its range individuals live up to about 1 year and spawn only once, whereas further north they reach over 2 years of age, spawning more than once. Like T. inermis, this species does not spawn in the extreme northern part of its range, so advection in the current systems is again important in maintaining the distribution. In the Pacific E. pacifica also shows this plastic character of changing maximum age with environmental variation. At the southern limit of its range individuals have a very short life span of only 6–8 months, whereas further north the maximum age is extended to about 15–21 months. In the most northern parts of the range the krill survive to over 2 years old and probably spawn twice. The maximum size across the range is about 20–22 mm, but growth is slower in the regions further north. As well as these general changes in development and life span, there are also sex-related differences. For example, in the northern Thysanossa species the males mature at just over 1 year old while the females mature mainly at over 2 years of age. In E. pacifica both sexes mature and spawn at 1 year old, but females may continue to survive and spawn at over 2 years old. In E. superba the situation can be different, with females spawning and maturing earlier at 2 years old, while males may not mature until over 3 years old. The euphausiids have the potential for rapid growth and development under suitable conditions, moulting as they increase in size. So, for example, E. superba has an energy input of perhaps 20% of body carbon per day or greater, sustained by a high and effective rate of filtration. This level of energy input can result in growth rates of 40.1 mm d 1,
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particularly for the younger age groups. Krill have the capacity for a large and sustained reproductive output under good conditions,with continuous or multiple spawning occurring through the season in some species. A key question for many of the species is how they survive during winter when food appears scarce, particularly in the extremely seasonal environments of the polar oceans. Studies of polar species show that the krill utilize stored lipids as a major energy source, but the dynamics of storage and utilization vary greatly between species. The lipids are accumulated primarily for winter survival or reproduction, but they may also provide a small degree of buoyancy that may help reduce the costs of swimming. In the Southern Ocean, the diet of E. superba varies with age. Phytoplankton sources are important for the early stages while older groups utilize more animal-based food sources or detritus. Lipids are utilized in winter, but a strong seasonal bloom of production is necessary for reproduction. For E. superba the suggestion is that winter survival is dependent not only on reduced metabolic rate, a potential reduction in size, and use of lipids, but also on the use of alternative food sources. Antarctic krill have been observed to get smaller during poor feeding conditions in the laboratory, but it is unclear how much this occurs in the ocean. Larval E. superba are dependent on sea ice as a habitat and the observation of krill grazing algae associated with the ice indicates that they can utilize this as an alternative food source. It remains unclear how important sea ice algae are for maintaining adult E. superba and this is likely to be a variable contribution to the diet depending on opportunity of access to the right feeding conditions. E. superba are also known to graze other components of the plankton, including copepods, so that a range of possible feeding strategies is likely to be open to them, depending on opportunity. E. crystallorophias occupies the area further south in the Antarctic where the spawning appears to occur before the main bloom, suggesting that lipid stores are used for survival and for reproduction. T. macrura has a similar distribution to E. superba but spawns earlier so it again is dependent on lipid stores for reproduction, but may also utilize other food sources to get through the winter. In northern regions T. inermis converts phytoplankton rapidly into lipids to cope with the seasonal environment, but also utilizes other available organic material such as detritus. M. norvegica also builds up high lipid stores but is more carnivorous, utilizing lipid-rich copepods.
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Overall, there appears to be a pattern from high to lower latitudes in the strategies of feeding and energy storage. Truly polar species such as E. crystallorophias and T. inermis rely totally on the seasonal phytoplankton bloom and lipid stores, whereas species such as E. superba, T. raschii and T. macrura survive winter utilizing alternative food sources and require the bloom to reproduce. Further away from the polar regions, species such as T. longicaudata and M. norvegica are more carnivorous, utilizing copepods as their main food source.
Spatial Distribution At large scales there are heterogeneities inthe distributions of euphausiids that extend for tens or hundreds of kilometers. Within these broad aggregations there are also more dense regions where krill form patches, swarms, or schools (Figure 4) forming a distribution generated by a very dynamic system, with aggregation and dispersal over a wide range of scales. These patches can be very dense and compact and it has been suggested that it is likely that on occasion all species of euphausiids aggregate to some extent. This ability to form such dense aggregations is certainly found in a number of species, particularly E. superba, but also
M. norvegica, Nyctyphanes australis, E. pacifica and E. lucens. The generation of such patchy distributions is the result of interactions between biological and physical processes over a range of scales. Over small scales and in very dense aggregations, behavior probably dominates. Swimming speeds can be high – B 20 cm s 1 in short bursts in Antarctic krill – so individuals have a marked ability to undertake directed movement at least over relatively short spatial scales. The formation of these smaller aggregations, along withdiurnal vertical migration, is considered to be mainly a predator avoidance effect. However, it will also lead to changes in the dynamics of the interaction of krill with their food and may generate complex outcomes in the dynamics of planktonic systems. At larger scales physical processes probably dominate so that aggregations are dependent on physical concentration mechanisms in areas of shelf-breaks, around islands, in ice edge regions or associated with eddies. This larger-scale aggregation may be a precondition for behavioral effects to dominate at smaller scales. Krill within aggregations appear to share more similar characteristics in size and maturity compared with those in other aggregations in the same area. The densities of krill within these aggregations can be well in excess of 10 000 m 3;
0.5 km
100 km
Figure 4 A hydroacoustic trace of an aggregation of Antarctic krill, E. superba.
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over 50 000 m 3 has been estimated for E. superba and E. pacifica and over 500 000 m 3 recorded for N. australis, E. lucens and M. norvegica. Some of the aggregations are extremely large and can account for a considerable proportion of the total biomass in an area. So for example, in one survey of Antarctic krill 410% of the regional biomass was recorded in just one aggregation that extended about 1 km horizontally. This large aggregation was observed in the vicinity of a large number of whales, suggesting that some of the very large, dense aggregations may be the result of intense predator–prey interaction and emphasizes the dynamic nature of the spatial distribution of krill. This makes the design of krill distribution surveys using nets or hydroacoustic techniques challenging and the survey data require careful interpretation and analysis. A number of the species undertake diurnal vertical migration, rising to nearer the surface and dispersing at night. This has been shown clearly in M. norvegica, whereas in E. superba vertical migration appears to be highly variable and may depend on local physical conditions, surface predator affects, and predation effects from below, particularly in areas of the shelf. In addition to the importance of predation effects the behavioral tracking of particular isolumes has been suggested as a mechanism involved in diurnal vertical migration and some species appear to show an endogenous rhythm. Seasonal changes in the pattern of aggregation and vertical migration have also been noted in some species. In one area it has been observed that during spring aggregations of E. superba are of the order of 0.7–2 km in length, whereas they are smaller and moredense in summer, and larger and less dense in autumn and winter. In the same study many of the swarms occurred in the upper 70 m during the summer, while in winter many were below 100 m deep. Other studies have found no such vertical change in depth distribution during the year, although diurnal vertical migration did change, being marked only during the spring and autumn.
Role in the Food Web Krill as Consumers
Krill show a range of feeding strategies from complete herbivory to total carnivory, with a full range of capabilities and flexible feeding strategies in between. In the Antarctic they can have a major impact on the large diatoms that form the major components of the intense blooms associated with the summer retreat of the sea ice. Antarctic krill are also known to consume copepods and negative
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correlations have been shown between the krill occurrence and the density of copepods and the phytoplankton concentration. This indicates that euphausiids are important in the plankton dynamics of these regions. They generate large fecal pellets which have high rates of sinking, suggesting that they can be important in the export of carbon from the surface layers. Rapid grazing of diatom blooms in some areas can therefore lead to a rapid flux of material to deeper ocean regions. The highly aggregated nature of the distributions is also likely to be important in determining the plankton dynamics, not just in terms of producing an interactive mosaic of production and consumption, but also in terms of the nutrient regime. Large krill swarms will generate high concentrations of ammonia that may favor the production of particular size groups of phytoplankton, leading to complex interactions in the plankton. Their role as consumers continues to be studied, but across the order they clearly have the capacity to feed on a wide range of food sources including diatoms, coccolithophores, dinoflagellates, chaetognaths, copepods, and other crustaceans; cannibalism has also been shown in some species. On the basis of observed variations in feeding strategies, it has been suggested that most species of euphausiids can adapt their feeding to utilize what is available, modifying their feeding strategies depending on the food they encounter. Krill as Prey
Krill are prey of many higher trophic level predators and as such play a key role throughout the oceans by transferring energy up the food chain. The baleen whales are the most well known predator, eating krill throughout their range, and despite the massive depletion of their populations due to harvesting they are still important krill predators. So for example, dense aggregations of M. norvegica and T. raschii in the Gulf of St Lawrence are associated with high abundances of fish (capelin) and a range of whale species including minkes, fin, blues, humpbacks, sperm, and beluga. Seals are also major predators of euphausiids in many areas. In Arctic waters, for example, harp seals consume M. norvegica and T. inermis, while in the Southern Ocean, crabeater seals consume E. crystallorophias. Further north, around some of the sub Antarctic islands such as South Georgia, the previously exploited fur seal populations that are now very large consume a considerable quantity of E. superba. Euphausiids also comprise a key component of the diet of a wide range of fish species, many of which
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are, or were, exploited. In the North Atlantic these include herring, cod, haddock, whiting, and mackerel. M. norvegica is probably the key species consumed, but others such as T. raschii, T. inermis and T. longicaudata are also important. In the North Pacific and adjacent regions E. pacifica is eaten by most commercial fish species, including Pacific cod, walleye pollack, chub mackerel, and sand lance. Other krill taken include T. raschii and T. inermis. The importance of euphausiids in the diet of many commercially exploited fish species is seen throughout the world. So for example around Australia N. australis is eaten by bluefin tuna and striped tuna, while in the Antarctic E. superba is eaten by the Mackerel icefish. Seabirds are also important predators of euphausiids throughout the world. In the North Atlantic a wide range of bird species consume M. norvegica and T. inermis including gulls, puffins, kittiwakes, and fulmars. The importance of sea-birds as predators of krill is highlighted in the Southern Ocean where E. superba is a key item in the diet and consumed in vast numbers by penguins (gentoos, macaronis, and Adelies) and by flying sea-birds including albatrosses such as grey-headed and black-browed albatross. This broad range view of euphausiids as prey emphasizes the important role that krill play in transferring energy to higher trophic levels in marine food webs worldwide. One of the key reasons for the importance of euphausiid species in food webs is the heterogeneity of krill distribution on a range of spatial and temporal scales. Different predators exploit the aggregation pattern with different foraging strategies, so exploiting different scales of pattern in the prey field. The pattern generated by the biological or biological–physical interactions will thus determine which predators can exploit the prey and hence the structure of the food web.
Krill Fisheries There are extensive fisheries for E. superba in the Southern Ocean, while in the Pacific off Japan there are important fisheries for E. pacifica. There is a more limited E. pacifica fishing off western Canada and there have also been intermittent fisheries for other species. The Southern Ocean fishery for E. superba is the largest and started at beginning of the 1970s, peaking in the early 1980s at 0.5 million tonnes. The fishery has since declined with changes in its economic basis. Catches over recent years have been o100 000 tonnes. The E. superba fishery in the Scotia Sea region is linked to seasonal seaice changes. During winter the fishery operates in the north
around South Georgia, it moves further south in the spring with the ice, to the area near the South Orkney Islands, and then during summer the fishery exploits krill around the Antarctic Peninsula. The management regime for E. superba takes account of krill recruitment variability, growth, and mortality to examine effects of various harvesting levels. Decision rules are included tomaintain stocks at a level that takes into account the dependent predators. At the current time, catch levels are much lower than the allowable catch (o10%) and future expansion of the fishery depends on the development of new products utilizing krill. An ecosystem approach is being developed for managing Southern Ocean fisheries and extensive predatormonitoring programs are operating. The challenge here is to develop management decision rules that consider not just the target species, the krill, but also incorporate ecological information from a number of levels in the food web,taking into account dependent species as well as environmental links. This ecosystem rather than species-based approach is one that will be increasingly relevant elsewhere.
Krill Variability There is considerable evidence of the importance of variation in the physical environment and circulation systems of the oceans in determining the distribution and abundance of krill. For example, links have been noted between variations in the oceanographic regimes associated with El Nin˜o events and the recruitment of E. pacifica in the North Pacific, while water temperature variations have been linked to 2–3 year variations in T. inermis populations. Biological processes associated with the environmental variation are also important in generating the variation observed in euphausiid populations. For example, there are marked interannual variations in the abundance of E. superba in the Southern Ocean, where recruitment strength has been linked to variations in the extent and concentration of sea ice. The current view is that increased sea ice cover and extent lead to favorable conditions for spawning and larval survival. The sea ice is thought to provide better over winter conditions for the krill. Salps compete with krill for phytoplankton – in poor sea ice years salp numbers are increased and krill recruitment is reduced. Further north in their range, E. superba abundance is dependent on the transport of krill in the ocean currents as well as fluctuations in the strength of particular cohorts. Given the importance of euphausiids in marine food webs throughout the world’s oceans, they are
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potentially important indicator species for detecting and understanding climate change effects. Changes in ocean circulation or environmental regimes will be reflected in changes in growth, development, recruitment success, and distribution. These effects may be most notable at the extremes of their distribution where any change in the pattern of variation will result in major changes in food web structure. Given their significance as prey to many commercially exploited species, this may also have a major impact on harvesting activities. A greater understanding of the large-scale biology of the euphausiids and the factors generating the observed variability is crucial. Obtaining good long-term and large-scale biological and physical data will be fundamental to this process.
See also Antarctic Circumpolar Current. Baleen Whales. Copepods. Marine Mammals: Sperm Whales and Beaked Whales.Phalaropes. Plankton. Sea Ice: Overview. Seals.
Further Reading
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Antarctic marine ecosystem: practical implementation of the Convention on the Conservation of the Antarctic Marine Living Resources (CCAMLR). ICES Journal of Marine Science 57: 778--791. Everson I (ed.) (2000) Krill: Biology, Ecology and Fisheries. Oxford: Blackwell Science. Everson I (2000) Introducing krill. In: Everson I (ed.) Krill: Biology, Ecology and Fisheries. Oxford: Blackwell Science. Falk-Petersen S, Hagen W, Kattner G, Clarke A, and Sargent J (2000) Lipids, trophic relationship, and biodiversity in Arctic and Antarctic krill. Canadian Journal of Fisheries and Aquatic Sciences 57: 178--191. Mauchline JR (1980) The biology of the Euphausids. Advances in Marine Biology 18: 371--677. Mauchline JR and Fisher LR (1969) The biology of the Euphausids. Advances in Marine Biology 7: 1--454. Miller D and Hampton I (1989) Biology and Ecology of the Antarctic Krill. BIOMASS Scientific Series, 9. Cambridge: SCAR & SCOR. Murphy EJ, Watkins JL, Reid K, et al. (1998) Interannual variability of the South Georgia marine ecosystem: physical and biological sources of variation. Fisheries Oceanography 7: 381--390. Siegel V and Nichol S (2000) Population parameters. In: Everson I (ed.) Krill: Biology, Ecology and Fisheries. Oxford: Blackwell Science.
Constable AJ, de la Mare W, Agnew DJ, Everson I, and Miller D (2000) Managing fisheries to conserve the
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KUROSHIO AND OYASHIO CURRENTS B. Qiu, University of Hawaii at Manoa, Hawaii, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1413–1425, & 2001, Elsevier Ltd.
Introduction The Kuroshio and Oyashio Currents are the western boundary currents in the wind-driven, subtropical and subarctic circulations of the North Pacific Ocean. Translated from Japanese, Kuroshio literally means black (‘kuro’) stream (‘shio’) owing to the blackish – ultramarine to cobalt blue – color of its water. The ‘blackness’ of the Kuroshio Current stems from the fact that the downwelling-dominant subtropical North Pacific Ocean is low in biological productivity and is devoid of detritus and other organic material in the surface water. The subarctic North Pacific Ocean, on the other hand, is dominated by upwelling. The upwelled, nutrient-rich water feeds the Oyashio from the north and leads to its nomenclature, parent (‘oya’) stream (‘shio’). The existence of a western boundary current to compensate for the interior Sverdrup flow is well understood from modern wind-driven ocean circulation theories. Individual western boundary currents, however, can differ greatly in their mean flow and variability characteristics due to different bottom topography, coastline geometry, and surface wind patterns that are involved. For example, the bimodal oscillation of the Kuroshio path south of Japan is a unique phenomenon detected in no other western boundary current of the world oceans. Similarly, interaction with the semi-enclosed and often ice-covered marginal seas and excessive precipitation over evaporation in the subarctic North Pacific Ocean make the Oyashio Current considerably different from its counterpart in the subarctic North Atlantic Ocean, the Labrador Current. Because the Kuroshio and Oyashio Current sexert a great influence on the fisheries, hydrography, and meteorology of countries surrounding the western North Pacific Ocean, they have been the focus of a great amount of observation and research in the past. This article will provide a brief review of the dynamic aspects of the observed Kuroshio and Oyashio Currents: their origins, their mean flow patterns, and their variability on seasonal-to-interannual timescales. The article consists of two sections, the first
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focusing on the Kuroshio Current and the second on the Oyashio Current. Due to the vast geographical areas passed by the Kuroshio Current (Figure 1), the first section is divided into three subsections: the region upstream of the Tokara Strait, the region south of Japan, and the Kuroshio Extension region east of the Izu Ridge. As will become clear, the Kuroshio Current exhibits distinct characteristics in each of these geographical locations owing to the differing governing physics.
The Kuroshio Current Region Upstream of the Tokara Strait
The Kuroshio Current originates east of the Philippine coast where the westward flowing North Equatorial Current (NEC) bifurcates into the northward-flowing Kuroshio Current and the southward-flowing Mindanao Current. At the sea surface, the NEC bifurcates nominally at 121N–131N, although this bifurcation latitude can change interannually from 111N to 14.51N. The NEC’s bifurcation tends to migrate to the north during El Nin˜o years and to the south during La Nin˜a years. Below the sea surface, the NEC’s bifurcation tends to shift northward with increasing depth. This tendency is due to the fact that the southern limb of the wind-driven subtropical gyre in the North Pacific shifts to the north with increasing depth. Branching northward from the NEC, the Kuroshio Current east of the Philippine coast has a mean geostrophic volume transport, referenced to 1250 dbar, of 25 Sv (1 Sverdrup ¼ 106 m3 s1). Seasonally, the Kuroshio transport at this upstream location has a maximum (B30 Sv) in spring and a minimum (B19 Sv) in fall. Similar seasonal cycles are also found in the Kuroshio’s transports in the East China Sea and across the Tokara Strait. As the Kuroshio Current flows northward passing the Philippine coast, it encounters the Luzon Strait that connects the South China Sea with the open Pacific Ocean (Figure 2). The Luzon Strait has a width of 350 km and is 2500 m deep at its deepest point. In winter, part of the Kuroshio water has been observed to intrude into the Luzon Strait and form a loop current in the northern South China Sea (see the dashed line in Figure 2). The loop current can reach as far west as 1171E, where it is blocked by the presence of the shallow shelf break off the south-east coast of China. The formation of the loop current is probably due to the north-east monsoon, prevailing
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KUROSHIO AND OYASHIO CURRENTS
60˚ N
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Figure 1 Schematic current patterns associated with the subtropical and subarctic gyres in the western North Pacific Ocean.
from November to March, which deflects the surface Kuroshio water into the northern South China Sea. During the summer months from May to September when the south-west monsoon prevails, the Kuroshio Current passes the Luzon Strait without intrusion. In the latitudinal band east of Taiwan (221N– 251N), the northward-flowing Kuroshio Current has been observed to be highly variable in recent years. Repeat hydrographic and moored current meter measurements between Taiwan and the southernmost Ryukyu island of Iriomote show that the variability of the Kuroshio path and transport here are dominated by fluctuations with a period of 100 days. These observed fluctuations are caused by impinging energetic cyclonic and anticyclonic eddies migrating from the east. The Subtropical Counter current (STCC) is found in the latitudinal band of 221N– 251N in the western North Pacific. The STCC, a shallow eastward-flowing current, is highly unstable due to its velocity shear with the underlying, westward-flowing NEC. The unstable waves generated by the instability of the STCC-NEC system tend to move westward while growing in amplitude. The
cyclonic and anticyclonic eddies that impinge upon the Kuroshio east of Taiwan are results of these large-amplitude unstable waves. Indeed, satellite measurements of the sea level (Figure 3) show that the Kuroshio east of Taiwan has higher eddy variability than either its upstream counterpart along the Philippine coast or its downstream continuation in the East China Sea. The Kuroshio Current enters the East China Sea through the passage between Taiwan and Iriomote Island. In the East China Sea, the Kuroshio path follows closely along the steep continental slope. Across the PN-line in the East China Sea (see Figure 2 for its location), repeat hydrographic surveys have been conducted on a quarterly basis by the Japan Meteorological Agency since the mid-1950s. Based on the measurements from 1955 to 1998, the volume transport of the Kuroshio across this section has a mean of 24.6 Sv and a seasonal cycle with 24.7 Sv in winter, 25.4 Sv in spring, 25.2 Sv in summer, and 22.8 Sv in fall, respectively. In addition to this seasonal signal, large transport changes on longer timescales are also detected across this section
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Figure 2 Schematic representation of the mean Kuroshio path (solid thick line) along the North Pacific western boundary. The thick dashed line south of Taiwan denotes the wintertime branching of the Kuroshio water into the Luzon Strait in the form of a loop current. PN-line denotes the repeat hydrographic section across which long-term Kuroshio volume transport is monitored (see Figure 4). Selective isobaths of 100 m, 200 m, and 1000 m are depicted.
(see Figure 4). One signal that stands out in the timeseries of Figure 4 is the one with the decadal timescale. Specifically, the Kuroshio transport prior to 1975 was low on average (22.5 Sv), whereas the mean transport value increased to 27.0 Sv after 1975. This decadal signal in the Kuroshio’s volume transport is associated with the decadal Sverdrup transport change in the subtropical North Pacific Ocean. Although the main body of the Kuroshio Current in the East China Sea is relatively stable due to the topographic constraint, large-amplitude meanders are frequently observed along the density front of the Kuroshio Current. The density front marks the shoreward edge of the Kuroshio Current and is located nominally along the 200 m isobath in the East China Sea. The frontal meanders commonly originate along the upstream Kuroshio front north east of Taiwan and they evolve rapidly while propagating downstreamward. The frontal meanders have typical wavelengths of 200–350 km, wave periods of 10–20 days, and downstreamward phase speeds of 10–25 cm s1. When reaching the Tokara Strait, the fully developed
frontal meanders can shift the path of the Kuroshio Current in the strait by as much as 100 km. Around 1281E–1291E and 301N, the Kuroshio Current detaches from the continental slope and veers to the east toward the Tokara Strait. Notice that this area is also where part of the Kuroshio water is observed to intermittently penetrate northward onto the continental shelf to feed the Tsushima Current. The frontal meanders of the Kuroshio described above are important for the mixing and water mass exchanges between the cold, fresh continental shelf water and the warm, saline Kuroshio water along the shelf break of the East China Sea. It is this mixture of the water that forms the origin of the Tsushima Current. The volume transport of the Tsushima Current is estimated at 2 Sv.
Region South of Japan
The Kuroshio Current enters the deep Shikoku Basin through the Tokara Strait. Combined surface current
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T/ P rms height variability (m) 60˚N
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Figure 3 Map of the root-mean-square (rms) sea surface height variability in the North Pacific Ocean, based on the TOPEX/ POSEIDON satellite altimetric measurements from October 1992 to December 1997. Maximum rms values of 40.4 m are found in the upstream Kuroshio Extension region south east of Japan. Sea surface height variability is also high in the latitudinal band east of Taiwan. (Adapted with permission from Qiu B (1999) Seasonal eddy field modulation of the North Pacific Subtropical Countercurrent: TOPEX/Poseidon observations and theory. Journal of Physical Oceanography 29: 2471–2486.)
35
Transport (Sv)
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Year Figure 4 Time-series of the geostrophic volume transport of the Kuroshio across the PN-line in the East China Sea (see Figure 2 for its location). Reference level is at 700 dbar. Quarterly available transport values have been low-pass filtered by the 1-year running mean averaging. (Data courtesy of Dr M. Kawabe of the University of Tokyo.)
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and hydrographic observations show that the Kuroshio’s volume transport through the Tokara Strait is about 30 Sv. Inference of transport from the sea level measurements suggests that the Kuroshio’s transport across the Tokara Strait is maximum in spring/summer and minimum in fall, a seasonal cycle similar to that found in the upstream Kuroshio Current. Further downstream, offshore of Shikoku, the volume transport of the Kuroshio has a mean value of 55 Sv. This transport increase of the Kuroshio in the deep Shikoku Basin is in part due to the presence of an anticyclonic recirculation gyre south of the Kuroshio. Subtracting the contribution from this recirculation reduces the mean eastward transport to 42 Sv. Notice that this ‘net’ eastward transport of the Kuroshio is still larger than its inflow transport through the Tokara Strait. This increased transport, B12 Sv, is probably supplied by the north-eastward-flowing current that has been occasionally observed along the eastern flank of the Ryukyu Islands. Near 1391E, the Kuroshio Current encounters the Izu Ridge. Due to the shallow northern section of the ridge, the Kuroshio Current exiting the Shikoku Basin is restricted to passing the Izu Ridge at either around 341N where there is a deep passage, or south of 331N where the ridge height drops. On interannual timescales, the Kuroshio Current south of Japan is known for its bimodal path fluctuations. The ‘straight path’, shown schematically by path A in Figure 5, denotes when the Kuroshio flows
closely along the Japan coast. The ‘large-meander path’, shown by path B in Figure 5, signifies when the Kuroshio takes a detouring offshore path. In addition to these two stable paths, the Kuroshio may take a third, relatively stable path that loops southward over the Izu Ridge. This path, depicted as path C in Figure 5, is commonly observed during transitions from a meandering state to a straight-path state. As the meander path of the Kuroshio can migrate spatially, a useful way of indexing the Kuroshio path is to use the mean distance of the Kuroshio axis from the Japan coast from 1321E to 1401E. South of Japan, the Kuroshio axis is well represented by the 161C isotherm. Based on this representation and seasonal water temperature measurements, Figure 6 shows the time-series of the Kuroshio path index from 1955 to 1998. A low index in Figure 6 denotes a straight path, and a high index denotes an offshore meandering path of the Kuroshio. From 1955 to 1998, the Kuroshio large meanders occurred in 1959–62, 1975–79, 1982–88, and 1990. Clearly, the large-meanders occurrence is aperiodic. Once formed, the meander state can persist over a period ranging from a year to a decade. In contrast, transitions between the meander and straight-path states are rapid, often completed over a period of several months. It is worth noting that development of the large meanders is often preceded by the appearance of a small meander south of Kyushu, which migrates eastward and becomes stationary after reaching 1361E.
36°N Honshu
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Figure 5 Schematic stable paths of the Kuroshio Current south of Japan. (Adapted with permission from Kawabe M (1985) Sea level variations at the Izu Islands and typical stable paths of the Kuroshio. Journal of the Oceanography Society of Japan 41: 307–326.) Selective isobaths of 1000 m, 2000 m, 4000 m, 6000 m, and 8000 m are included.
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2.5
Distance (°lat)
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1.5
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0.5 1955
1960
1965
1970
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1980
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1995
Year Figure 6 Time-series of the Kuroshio path index from 1955 to 1998, where the Kuroshio path index is defined as the offshore distance of the Kuroshio axis (inferred from the 161C isotherm at the 200 m depth) averaged from 1321 to 1401E. Solid dots denote the seasonal index values and the solid line indicates the annual average. (Adapted with permission from Qiu B and Miao W (2000) Kuroshio path variations south of Japan: Bimodality as a self-sustained internal oscillation. Journal of Physical Oceanography 30: 2124–2137.)
Several mechanisms have been proposed to explain the bimodal path variability of the Kuroshio south of Japan. Most studies have examined the relationship between the Kuroshio’s path pattern and the changes in magnitude of the Kuroshio’s upstream transport. Earlier studies of the Kuroshio path bimodality interpreted the meandering path as stationary Rossby lee wave generated by the protruding coastline of Kyushu. With this interpretation, the Kuroshio takes a meander path when the upstream transport is small and a straight path when it is large. By taking into account the realistic inclination of the Japan coast from due east, more recent studies have provided the following explanation. When the upstream transport is small, the straight path is stable as a result of the planetary vorticity acquired by the north-eastwardflowing Kuroshio being balanced by the eddy dissipation along the coast. When the upstream transport is large, positive vorticity is excessively generated along the Japan coast, inducing the meander path to develop downstream. In the intermediate transport range, the Kuroshio is in a multiple equilibrium state in which the meandering and straight paths coexist. Transitions between the two paths in this case are determined by changes in the upstream transport (e.g. the transition from a straight path to a meander path requires an increase in upstream transport). A comparison between the Kuroshio path variation (Figure 6) and the Kuroshio’s transport in the upstream East China Sea (Figure 4) shows that the 1959–62 large-meander event does correspond to a large upstream transport. However, this correspondence becomes less obvious after 1975, as there were times when the upstream transport was large, but no large meander was present. Assuming
that the upstream Kuroshio transport after 1975 is in the multiple equilibrium regime, the correspondence between the path transition and the temporal change in the upstream transport (e.g. the required transport increase for the transition from a straight path to a meander path) is also inconclusive from the timeseries presented in Figures 4 and 6. Given the low frequency and irregular nature of the Kuroshio path changes, future studies based on longer transport measurements are needed to further clarify the physics underlying the Kuroshio path bimodality. Downstream Extension Region
After separating from the Japan coast at 1401E and 351N, the Kuroshio enters the open basin of the North Pacific Ocean where it is renamed the Kuroshio Extension. Free from the constraint of coastal boundaries, the Kuroshio Extension has been observed to be an eastward-flowing inertial jet accompanied by large-amplitude meanders and energetic pinched-off eddies. Figure 7 shows the mean temperature map at 300 m depth, in which the axis of the Kuroshio Extension is well represented by the 121C isotherm. An interesting feature of the Kuroshio Extension east of Japan is the existence of two quasi-stationary meanders with their ridges located at 1441E and 1501E, respectively. The presence of these meanders along the mean path of the Kuroshio Extension has been interpreted as standing Rossby lee waves generated by the presence of the Izu Ridge. A competing theory also exists that regards the quasi-stationary meanders as being steered by the eddy-driven abyssal mean flows resulting from instability of the Kuroshio Extension jet.
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2 4
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5 6
4 40˚
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Sea of Japan
8
Latitude
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9 10
10 Japan
11 16
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15
15 140˚
150˚
160˚ Longitude
14 170˚
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Figure 7 Mean temperature map (1C) at the 300 m depth from 1976 to 1980. (Adapted with permission from Mizuno K and White WB (1983) Annual and interannual variability in the Kuroshio Current System. Journal of Physical Oceanography 13: 1847–1867.)
Near 1591E, the Kuroshio Extension encounters the Shatsky Rise where it often bifurcates. The main body of the Kuroshio Extension continues eastward, and a secondary branch tends to extend northeastward to 401N, where it joins the eastwardmoving Subarctic Current. After overriding the Emperor Seamounts along 1701E, the mean path of the Kuroshio Extension becomes broadened and instantaneous flow patterns often show a multiplejet structure associated with the eastward-flowing Kuroshio Extension. East of the dateline, the distinction between the Kuroshio Extension and the Subarctic Current is no longer clear, and together they form the broad, eastward-moving North Pacific Current. As demonstrated in Figure 3, the Kuroshio Extension region has the highest level of eddy variability in the North Pacific Ocean. From the viewpoint of wind-driven ocean circulation, this high eddy variability is to be expected. Being a return flow compensating for the wind-driven subtropical interior circulation, the Kuroshio originates at a southern latitude where the ambient potential vorticity (PV) is relatively low. For the Kuroshio to smoothly rejoin the Sverdrup interior flow at the higher latitude, the low PV acquired by the Kuroshio in the south has to be removed by either dissipative or nonlinear forces along its western boundary path. For the narrow and swift Kuroshio Current, the dissipative force is insufficient to remove the low PV anomalies. The consequence of the Kuroshio’s inability to effectively diffuse the PV anomalies along its path results in the accumulation of low PV water in its extension region, which generates an anticyclonic recirculation gyre and provides an energy source for flow instability. Due
to the presence of the recirculation gyre (Figure 8), the eastward volume transport of the Kuroshio Extension can reach as high as 130 Sv south east of Japan. This is more than twice the maximum Sverdrup transport of about 50 Sv in the subtropical North Pacific. The inflated eastward transport is due to the presence of the recirculating flow to the south of the Kuroshio Extension. Although weak in surface velocity, Figure 8 shows that the recirculating flow has a strong barotropic (i.e. depthindependent) component. As a consequence, the volume transport of the recirculation gyre in this case is as large as 80 Sv. In addition to the high meso-scale eddy variability, the Kuroshio Extension also exhibits large-scale changes on interannual timescales. Figure 9A and B compares the sea surface height field in the Kuroshio Extension region in November 1992 with that in November 1995. In 1992, the Kuroshio Extension had a coherent zonal-jet structure extending beyond the dateline. The zonal mean axis position of the Kuroshio Extension from 1411E to 1801E in this case was located north of 351N. In contrast, the jet-like structure in 1995 was no longer obvious near 1601E and the zonal mean axis position shifted to 341N. Note that the changes in the zonal mean axis position of the Kuroshio Extension have interannual timescales (Figure 9C) and are associated with the changes in the strength of the southern recirculation gyre. As the recirculation gyre intensifies (as in 1992), it elongates zonally, increasing the zonal mean eastward transport of the Kuroshio Extension and shifting its mean position northward. When the recirculation gyre weakens (as in 1995), it decreases the eastward transport of the Kuroshio Extension and shifts its zonal mean position southward. At
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KUROSHIO AND OYASHIO CURRENTS
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V: (32 CW) 145 _141° E (WHP P10) 6 _ 9 Nov. 1993 0
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Figure 8 North-eastward velocity profile from lowered acoustic Doppler current meter profiler (ADCP) measurements along the WOCE P10 line south east of Japan in November 1993 (see the dashed line in Figure 7 for its location). Units are cm s1 and southwestward flow is shaded. (Figure courtesy of Drs E. Firing and P. Hacker of the University of Hawaii.)
present, the cause of the low-frequency changes of the recirculation gyre is unclear.
The Oyashio Current Due to the southward protrusion of the Aleutian Islands, the wind-driven subarctic circulation in the North Pacific Ocean can be largely divided into two cyclonic subgyres: the Alaska Gyre to the east of the dateline and the Western Subarctic Gyre to the west (Figure 1). To the north, these two subgyres are connected by the Alaskan Stream, which flows southwestward along the Aleutian Islands as the western boundary current of the Alaska Gyre. Near the dateline, the baroclinic volume transport of the Alaskan Stream in the upper 3000 m layer is estimated at about15–20 Sv. As the Alaskan Stream flows further westward, the deep passages between 1681E and 1721E along the western Aleutian Islands allow part of the Alaskan Stream to enter the Bering Sea. In the deep part of the Bering Sea, the intruding Alaskan Stream circulates anticlockwise and forms
the Bering Sea Gyre. The western limb of the Bering Sea Gyre becomes the East Kamchatka Current, which flows south-westward along the east coast of the Kamchatka Peninsula. The remaining part of the Alaskan Stream continues westward along the southern side of the Aleutian Islands and upon reaching the Kamchatka Peninsula, it joins the East Kamchatka Current as the latter exits the Bering Sea. As the East Kamchatka Current continues southwestward and passes along the northern Kuril Islands, some of its water permeates into the Sea of Okhotsk. Inside the deep Kuril Basin in the Sea of Okhotsk, the intruding East Kamchatka Current water circulates in a cyclonic gyre. Much of this intruding water moves out of the Sea of Okhotsk through the Bussol Strait (46.51N, 151.51E), where it joins the rest of the south-westward-flowing East Kamchatka Current. The East Kamchatka Current is renamed the Oyashio Current south of the Bussol Strait. Because of the intrusion in the Sea of Okhotsk, the water properties of the Oyashio Current are different from those in the upstream East
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40°N
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33°N 1986 1987 1988 1989 1990 1991 1992 1993 1994 1995 1996 1997 1998 1999 (C) Year Figure 9 Sea surface height maps on (A) 20 November 1992 and (B) 15 November 1995 from the TOPEX/POSEIDON altimeter measurements. (C) Time-series of the mean axis position of the Kuroshio Extension from 1411E to 1801. (Adapted with permission from Qiu B (2000) Interannual variability of the Kuroshio Extension system and its impact on the wintertime SST field. Journal of Physical Oceanography 30: 1486–1502.)
Kamchatka Current. For example, the mesothermal water present in the East Kamchatka Current (i.e. the subsurface maximum temperature water appearing in the halocline at a depth of 150–200 m) is no longer observable in the Oyashio. While high dissolved oxygen content is confined to above the halocline in the upstream East Kamchatka Current, elevated dissolved oxygen values can be found throughout the upper 700 m depth of the Oyashio water. The baroclinic volume transport of the Oyashio Current along the southern Kuril Islands and off Hokkaido has been estimated at 5–10 Sv from the geostrophic calculation with a reference level of nomotion at 1000 or 1500 m. Combining moored
current meter and CTD (conductivity-temperaturedepth) measurements, more recent observations along the continental slope south east of Hokkaido show that the Oyashio Current has a well-defined annual cycle: the flow tends to be strong, reaching from surface to bottom, in winter/spring, and it is weaker and confined to the layer shallower than 2000 m in summer and fall. The total (baroclinic þ barotropic) volume transport reaches 20–30 Sv in winter and spring, whereas it is only 3–4 Sv in summer and fall. This annual signal in the Oyashio’s total transport is in agreement with the nnual signal in the Sverdrup transport of the wind-driven North Pacific subArctic gyre.
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KUROSHIO AND OYASHIO CURRENTS
120˚E
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20˚N Figure 10 Water temperature map at the 100 m depth in September 1989 compiled by the Japan Meteorological Agency. Contour interval is 11C.
After flowing south-westward along the coast of Hokkaido, the Oyashio Current splits into two paths. One path veers offshoreward and contributes to the east-north-eastward-flowing SubArctic Current. This path can be recognized in Figure 10 by the eastward-veering isotherms along 421N south east of Hokkaido. Because the Oyashio Current brings water of subarctic origin southward, the SubArctic Current is accompanied by a distinct temperaturesalinity front between cold, fresher water to the north and warm, saltier water of subtropical origin to the south. This water mass front, referred to as the
Oyashio Front or the Subarctic Front, has indicative temperature and salinity values of 51C and 33.8 PSU at the 100 m depth. Across 1651E, combined moored current meter and CTD measurements show that the SubArctic Current around 411N has a volume transport of 22 Sv in the upper 1000 m layer. The second path of the Oyashio Current continues southward along the east coast of Honshu and is commonly known as the first Oyashio intrusion. As shown in Figure 10, an addition to this primary intrusion along the coast of Honshu, the southerly Oyashio intrusion is also frequently observed further
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Figure 12 (A) Spatial pattern of the second empirical orthogonal function (EOF) mode of the wintertime sea surface temperature anomalies (1950–1992) in the Pacific Ocean. This mode explains 11% of the variance over the domain. (B) Time-series of the wintertime sea surface temperature anomalies averaged in the Kuroshio-Oyashio outflow region (321N–461N,1361E–1761W). (Adapted with permission from Deser C and Blackmon ML (1995) On the relationship between tropical and North Pacific sea surface variations. Journal of Climate 8: 1677–1680.)
offshore along 1471E. This offshore branch is commonly known as the second Oyashio intrusion. The annual mean first Oyashio intrusion east of Honshu reaches on average the latitude 38.71N, although in some years it can penetrate as far south as 371N (see Figure 11). In addition to the year-to-year fluctuations, Figure 11 shows that there is a trend for the Oyashio Current to penetrate farther southward after the mid-1970s. Both this long-term trend and the interannual changes in the Oyashio’s intrusions seem to be related to the changes in the intensity of the Aleutian low atmospheric pressure system and the southward shift in the position of the mid-
latitude westerlies. It is worth noting that the anomalous southward intrusion of the Oyashio Current not only influences the hydrographic conditions east of Honshu, it also affects the environmental conditions in the fishing ground and the regional climate (e.g. an anomalous southward intrusion tends to decrease the air temperature over eastern Japan).
Concluding Remarks Because the Kuroshio and Oyashio Currents transport large amounts of water and heat efficiently
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KUROSHIO AND OYASHIO CURRENTS
in the meridional direction, there has been heightened interest in recent years in understanding the dynamic roles played by the time-varying Kuroshio and Oyashio Currents in influencing the climate through sea surface temperature (SST) anomalies. Indeed, outside the eastern equatorial Pacific Ocean, the largest SST variability on the interannual-todecadal time-scale in the North Pacific Ocean resides in the Kuroshio Extension and the Oyashio outflow regions (Figure 12). Large-scale changes in the Kuroshio and Oyashio current systems can affect the SST anomaly field through warm/cold water advection, upwelling through the base of the mixed layer, and changes in the current paths and the level of the meso-scale eddy variability. At present, the relative roles played by these various physical processes are not clear. This article summarizes many observed aspects of the Kuroshio and Oyashio Current systems, although due to the constraints of space, important subjects such as the water mass transformation processes in regions surrounding the Kuroshio and Oyashio and the impact of the Kuroshio and Oyashio variability upon the oceanographic conditions in coastal and marginal sea areas have not been addressed. It is worth emphasizing that our knowledge of the Kuroshio and Oyashio Currents has increased significantly due to the recent World Ocean Circulation Experiment (WOCE) program (observational phase: 1990–1997). Fortunately, many of the observational programs initiated under the WOCE program are being continued. With results from these new observations, we can expect an improved description of the Kuroshio and Oyashio Current systems in the near future, especially of the variability with timescales longer than those described in this article.
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Further Reading Dodimead AJ, Favorite JF, and Hirano T (1963) Review of oceanography of the subarctic Pacific region. Bulletin of International North Pacific Fisheries Commission 13: 1--195. Kawabe M (1995) Variations of current path, velocity, and volume transport of the Kuroshio in relation with the large meander. Journal of Physical Oceanography 25: 3103--3117. Kawai H (1972) Hydrography of the Kuroshio Extension. In: Stommel H and Yoshida K (eds.) Kuroshio – Its Physical Aspects, pp. 235--354. Tokyo: University of Tokyo Press. Mizuno K and White WB (1983) Annual and interannual variability in the Kuroshio Current system. Journal of Physical Oceanography 13: 1847--1867. Nitani H (1972) Beginning of the Kuroshio. In: Stommel H and Yoshida K (eds.) Kuroshio – Its Physical Aspects, pp. 129--163. Tokyo: University of Tokyo Press. Pickard GL and Emery WJ (eds.) (1990) Descriptive Physical Oceanography: An Introduction, 5th edn. Oxford: Pergamon Press. Shoji D (1972) Time variation of the Kuroshio south of Japan. In: Stommel H and Yoshida K (eds.) Kuroshio – Its Physical Aspects, pp. 217--234. Tokyo: University of Tokyo Press. Taft BA (1972) Characteristics of the flow of the Kuroshio south of Japan. In: Stommel H and Yoshida K (eds.) Kuroshio – Its Physical Aspects, pp. 165--216. Tokyo: University of Tokyo Press. Tomczak M and Godfrey JS (1994) Regional Oceanography: An Introduction. Oxford: Pergamon Press.
See also Abyssal Currents. Okhotsk Sea Circulation. Pacific Ocean Equatorial Currents. Wind Driven Circulation.
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LABORATORY STUDIES OF TURBULENT MIXING J. A. Whitehead, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction As described elsewhere in this encyclopedia, turbulence and diffusion act both vertically in mixing up the density field of the ocean and laterally in mixing adjacent density and tracer fields. Vertical mixing of density in a stratified fluid dissipates turbulent kinetic energy by raising the potential energy of the density field, and the ratio of this dissipation to viscous dissipation of turbulence is a fundamental quantity needed to understand the energy balance of the ocean. Estimates of these rates by theory are not yet readily available in a form useful for the ocean. Numerical approaches and computational fluid dynamics are under development, but they are not capable of investigating a wide range of parameter space in present form, nor can results be recovered that are verified by experimental benchmarks. Direct ocean measurements rely on theoretical assumptions about the form of turbulence, which must ultimately be verified in the laboratory. Therefore, laboratory measurements continue to be essential to the quest of determining the rates of turbulent dissipation within flows of stratified fluid. Both the production of turbulence through instability and its dissipation are markedly different from the case of instability and dissipation within a homogeneous fluid. Ignoring internal wave radiation from turbulent regions, dissipation is partitioned between viscous dissipation and the work that increases potential energy. This is shown below by the two integrals for energy of a simple system with no internal body forces and closed bottom and top boundaries in a field of gravity. In this example, the flow must start with an initial value of kinetic energy, and the decrease in kinetic energy of a volume of incompressible fluid is equal to the rate of buoyancy work plus viscous dissipation: D E 1 dhv˜ v˜ i ¼ ghrwi v ðrv˜ Þ2 2 dt The change in potential energy is the rate of buoyancy work plus buoyancy flux multiplied by elevation: Z h dhrzi z½rw dz ¼ ghrwi þ g g dt 0
where angle brackets are averages over all three spatial dimensions, and the square brackets are over the two lateral dimensions. Here, v˜ is the threedimensional velocity vector, n is viscous diffusivity, the variables z, w are the vertical direction and velocity (the direction of gravity g), and density is r. Clearly, if the following three conditions are met, then the only two terms that can dissipate the kinetic energy are viscous dissipation and buoyancy (heat) flux: first, the potential energy is not changing in time; second, dense (considering it to be cold) fluid enters the bottom with lighter (warm) fluid leaving the top (it would require a downward heat flow into the volume to allow this); and third, the volume is given an initial value of kinetic energy that is allowed to run down. It is the purpose of laboratory experiments to allow measurements of density and velocity fields and to obtain the partition between viscous dissipation and buoyancy flux. No instrument exists for precisely measuring every term within any of the above brackets, so simplified approaches have been necessitated.
Experiments Two types of experiments generate the turbulence, either a shear-flow instability is set up or eddies are directly generated. In addition, there are two groups of density distribution, one with sharp interfaces and the other having continuous stratification. Since both of the dissipation terms shown above are negative, there are no cases with both of the equations in their steady form. Experiments incorporate either transient setups that run down with time, or utilize flowing tanks (mostly with salt-stratified water but a few wind tunnels with thermal stratification are used too). In the latter case, the flows are transient following a fluid parcel. The techniques to produce eddies are numerous. Figure 1 shows some of the laboratory configurations. Some generate turbulence behind grids in tunnels (Figure 1(a)) and others are driven by buoyancy (Figure 1(b)). In some flumes and wind tunnels, narrowing the sides enhances the shear in a test region. Other experiments have a moving grid or rod stirrer (Figure 1(c)), and still others have a moving lid (Figure 1(d)). Not sketched are special studies of pumped jets directed toward an interface, and experiments with double-diffusion driven flows, both being directed toward explicit mechanisms of mixing. Experiments have been motivated by numerous phenomena in addition to oceanographic ones; some
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Figure 1 Configurations of various laboratory experiments to measure stratified mixing.
examples are fire prevention, ventilation, snow avalanches, ecosystem studies, solar ponds, mixing of industrial chemicals, and turbidity currents. Flows have been produced in pumped tunnels, as sketched in Figure 1(a), with the stratification produced by either temperature (using heat-transfer devices) or salinity (using multiple sources). A variation is a closed-circuit tunnel with a special device for turbulence-free propulsion of the water. In some cases the vertical salinity distribution is initially set and the density profile runs down with time. Another variant has layers pumped in at multiple elevations with the same amount withdrawn at each corresponding level downstream and returned to cisterns. Also, there are currents driven by buoyant flow in tanks with sloping tops and bottoms, as shown in Figure 1(b), or with exchange flows in passages between reservoirs containing waters of differing salinity. Then, there are experiments in closed containers with oscillating grids or moving rods as sketched in Figure 1(c). In some cases the experiments are in annular chambers and some are rotating on a turntable (Figure 1(d)). Salt-stratified experiments are the most numerous. They possess a vertical salinity distribution that evolves with time, although
there are also thermal experiments using air or water motivated by engineering applications. The dynamics are all characterized by velocity scale of the turbulence u (which may be the same size as velocity difference in a sheared laminar flow whose instability generates the turbulence) and density difference Dr. The force of gravity makes the density difference equivalent to a buoyancy difference g0 ¼ gDr/r0. Geometrically, there is the separation distance d between regions of different velocity and density. Finally, there are two additional fluid properties, the viscosity n and density diffusivity D. Velocity, reduced gravity, length, viscosity, and diffusivity can be reduced to many combinations of three dimensionless numbers. For stratified mixing, they are picked sequentially to represent the important balances in the flow in rank order. The primary dimensionless number is the bulk Richardson number Ri ¼ g0 d=u2. It is a measure of the ratio of buoyancy to fluid inertia. The second number is the Reynolds number Re ¼ u d=v, which is the ratio of inertial to viscous force. The third is the Schmidt number Sc ¼ n/D. It is the ratio of viscosity to density diffusivity. Experiments generally have the objective to determine a buoyancy work rate (frequently called buoyancy flux) as a function of these three numbers. By the early 1990s it was clear that the buoyancy flux obeyed a range of power-law relations with Ri, and that these are sensitive to details of actual experiments in most cases. The relative roles of Re and Sc are less well documented. Virtually any source of turbulence can be used to mix stratified fluid, and one of the challenges of experiments is to separate the influence of the spatial distribution of the turbulent source from the actual processes within the fluid. Turbulence that is shed from a vertical rod that moves laterally and sheds a turbulent wake in fresh water above a salt layer (Figure 2(a)) produces striations that are strongest near the interface and weaker at higher and lower levels. This produces a divergence in buoyancy flux so the interface gets progressively thicker with time. However, continuous stratification (Figure 2(b)) can spontaneously break down to internal layers. Therefore, in both cases, the flux varies locally. To complicate matters, the variation of the stratification is usually about the same size as the scale of the turbulence, so in almost all experiments it is not obvious that statistical turbulence theory applies. Probably the simplest configuration has salt water under fresh water with grid stirring confined to one layer, for example, the top layer. This easily attains high Reynolds numbers using a horizontal grid moving up and down with vigorous oscillatory motion. To determine u at the level of the interface,
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LABORATORY STUDIES OF TURBULENT MIXING
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10−5 Figure 2 Shadowgraphs from a parallel beam of light falling onto a screen. They show the effect of turbulence produced by many excursions of a moving rod (a) for a layered fluid as in Figure 1(c) right (salt water under fresh, the rod has recently reversed direction near the tank wall); and (b) for a stratified fluid that breaks down into layers as in Figure 1(c) left (stratified salt water, the rod is moving toward the left into a placid, previously mixed fluid).
grid velocity should be multiplied by a suitable constant to account for spatial variation of the turbulence between grid and interface. As time progresses, for Ri41 the turbulence in the upper layer causes the interface to remain sharp and it mixes salt water up into the top layer. This ‘entrains’ salt water into the top layer, which increases both the volume and salinity of the top layer and decreases the volume of the bottom one but leaves its salinity unchanged. The interface moves downward with entrainment velocity ue, and the speed quantifies the mixing rate. As density difference between the layers decreases, Ri decreases. The entrainment velocity increases steeply with decreasing Richardson number as shown by solid circles in Figure 3. For Rio1, the interface deflection is as large as d and the subsequent mixing rate is rapid and soon the two layers mix completely. Many other experiments have produced entrainment velocity measurements; a collection of some is shown in the lower cluster of Figure 3. In some of them, the mixing is supplied by shear instability driven by a rotating screen in an annulus with stratified fluid (Figure 1(d)). A mixed layer with a sharp density jump at the bottom penetrates into the stratified fluid. Others involve gravity currents (Figure 1(b)), with buoyant outflows and with counter-flows. All of these configurations successfully give useful data for large stratification (Rio1). The scatter in the points shown in Figure 3 is typical and is due to the statistical nature of the data rather
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Ri Figure 3 Entrainment velocity as a function of Richardson number for assorted experiments. Solid circles: Grid experiments where the grid velocity is to define u * . Open circles: Buoyant outflows. Triangles: Density currents. Crosses: Counterflows. Data and slopes taken from Fernando HJS (1991) Turbulent mixing in stratified fluids. Annual Review of Fluid Mechanics 23: 455–493, figure 15, with permission from author, and Turner JS (1973) Buoyant convection from isolated sources. Buoyancy Effects in Fluids, pp. 165–206, Figure 9.3. Cambridge, UK: Cambridge University Press, with permission from author.
than instrumental error. In such experiments, the horizontally averaged density typically breaks up into patches and layers so that local regions have different local values of Richardson number. In spite of such scatter, the results of such experiments are overwhelmingly consistent with each other with respect to the general trends of the data shown in Figure 3. Primarily, the Richardson number is the most important variable governing mixing rate if it is of order 1 or more. The Reynolds number does exhibit some role especially if less than c. 500. Experiments to date range up to almost Reo105 and generally speaking the mixing is sensitive to Reynolds number for the entire range. It is thought that mixing will finally become insensitive for very large Re but data up to such a possible limit are not yet available. Three power laws are sketched as straight lines in this figure. To the left the open circle data have a smaller slope. To the right, the data are inversely proportional to a higher power of Richardson number. The constants of proportionality are functions of the actual configuration and the manner of defining velocity and density. As an example, the
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relation found by Kato and Phillips for shear-driven experiments is
10−1
ue ¼ 2:5 Ri1 u This can be rearranged to gDrue d ¼ 2:5r0 u3 , which states that the rate of change of potential energy is proportional to the rate of delivery of kinetic energy for mixing. The value of power-law dependence has been extensively studied and discussed. There are some circumstances in which there is no dependence because stratification effects are smaller than viscous, diffusive, or turbulent effects. In other cases there are over 40 proposed relations with Ri, but there is still no clear consensus about the range of validity for each of these in Ri, Re, and Sc space. This lack of agreement seems to arise for a number of reasons. First, no tight cluster about one line is found because of the scatter mentioned above. Second, it has always been found that the results tend to be specific for each experimental configuration. Third, the experiments are in water or air, so only a few values of Sc are investigated. One proposed relation between the three dimensionless numbers has the dimensionless entrainment at successively increasing Ri proportional to Ri 3/2, ScRi 1, and Sc 1/2 Re 1/4. It shows that there is an increasingly important role for molecular and turbulence effects as stratification is increased. However, in listing 30 such relations, Fernando found that the 1.5 power law generally tended to be for higher values of Ri rather than lower values. A gravity current down a slope is of particular interest to oceanography because of its relevance to deep overflows in polar regions and salt plumes in aridpregions. In such studies the Froude number Fr ¼ ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi u = g0 d cos yðERi1=2 Þ (y is angle of the slope) is often used to measure the intensity of mixing compared to stratification. Results for a wide range of gravity-driven currents are shown in Figure 4. As in Figure 3 it is clear that large stratification suppresses entrainment velocity, and that ueEu with smaller stratification. In addition, it is clear that mixing is enhanced at larger Reynolds numbers.
Continuous Stratification The layers of mixed fluid in mechanical stirring experiments are separated by very sharp interfaces, so for large values of Ri the layers remain well defined. Therefore, the results are relatively precise and easy to interpret. Of course, experiments with continuous stratification are more similar to the ocean. Continuously stratified fluid exposed to turbulence tends
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Fr = u∗ /(g ′dcos())1/2 Figure 4 Entrainment velocity in laboratory gravity currents down a slope compared to estimates of ocean overflows. The solid triangle is for Lake Ogawara, the solid square is for Mediterranean outflow into the Atlantic, the small star is for the Denmark Strait overflow, and the solid diamond is for the Faroe Bank channel overflow. For rotating density currents, the open squares, open triangles, and large stars are experiments with Reo100 and the open diamonds are with ReZ100. The shaded area and open circles are found for large Reynolds number nonrotating density currents. Supplied by C. Cenedese.
to break up into layers for large Ri. Therefore, N varies locally and hence the dimensionless number varies locally. As a result, the mixing becomes concentrated in local regions. In addition, as time progresses the layers evolve and slowly change flux. The dynamics that determine the size of the layers remain controversial. In some cases, the layer depth scales with the scale u/N, and in other cases the scales are linked to vortical modes shed from the stirrers, or even to the stirrer size itself. In continuously stratified experiments, the overall Richardson number can be defined as Ric ¼ (Nd/u )2 or, if shear du/dz is imposed, as Ris ¼ (N/(du/dz))2. The invention of the bathythermograph led to the discovery of extensive layering within the ocean. Although this layering can be explained as a consequence of localized wave breaking, its universal character suggests that there is a more fundamental cause. Laboratory experiments with stirring near the sidewall of a stratified fluid, and later experiments with continuously stratified fluid stirred with a rod, all exhibited the spontaneous growth of layers for Ricc1. Figure 5 shows this growth in salt-stratified water with grid-generated turbulence (with d set to the mesh size) for Ric ¼ 10.7. In some elevations the local stratification (as measured by N) increases, and
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at other elevations it decreases toward zero. This has a profound influence on propagation of both internal and acoustic waves through the water. As time progresses beyond the stage shown in the figure, the layers begin to interact with each other and some will be eliminated, so finally two layers remain with only one interface in between. Ultimately, the density difference between these two layers decreases to zero, and the fluid becomes fully mixed. In contrast, such an experiment with values of Ric approximately 1 or lower does not produce layers (Figure 6). Instead, a mixed layer forms at the bottom and top of the fluid. Simultaneously, the interior stratification gradually decreases. The result is a fluid with values of N decreasing everywhere. Figures 5 and 6 were produced with data from experiments with the configurations shown in Figure 1(c). Accurate resolution of the vertical density field by a conductivity microprobe (developed for stratified turbulent flume measurements) allows precise measurements of the change of potential energy with time. This change is quantified by a flux Richardson number Rfc, defined as rate of change in potential energy divided by power (rate of energy) exerted by the stirrer (which is estimated for the grid using known drag laws). The resulting data are shown in Figure 7. Starting from Ric ¼ 0, experiments with increasing values have increasing values of Rfc, which level off at a value RfcC0.067 at RicC1. Almost all layered and continuously stratified experiments can be interpreted as having a flux Richardson number that reaches its maximum value
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(kg m−3) Figure 6 The evolution of a density field with stirring source throughout the entire fluid and with Rir1. Adapted from Rehmann CR and Koseff JR (2004) Mean potential energy change in stratified grid turbulence. Dynamics of Atmospheres and Ocean 37: 271–294, with permission from Elsevier.
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Figure 7 Mixing efficiency in grid-stirring experiments with continuous stratification and small Ri. Adapted from Rehmann CR and Koseff JR (2004) Mean potential energy change in stratified grid turbulence. Dynamics of Atmospheres and Ocean 37: 271–294, with permission from Elsevier.
of c. 0.1 at a Richardson number of order 1, with values decreasing toward 0 for larger and smaller values. The exact value of the maximum has been widely discussed, with some estimates approaching a maximum value of 0.2 and others only reaching a maximum value of 0.05 or so. Naturally, the exact definition depends on the choice of a length and velocity scale, which is not only a matter of choice, but also subjected to the details of each apparatus and analysis technique. In addition, since most experiments are either transient or possessing a variation in space, the measurement location
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Rfc
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Figure 8 Mixing efficiency as a function of Reynolds number in mixing grid experiments of a uniformly stratified fluid. Stratification was from salinity in some experiments, temperature in others, and from both in one case. Adapted from Rehmann CR and Koseff JR (2004) Mean potential energy change in stratified grid turbulence. Dynamics of Atmospheres and Ocean 37: 271–294, with permission from Elsevier.
In summary, laboratory experiments to measure the rate of mixing of stratified fluids by turbulence have been conducted using a wide range of devices. Both qualitative and quantitative data provide information to oceanographers and numerical modelers. To estimate the mixing rates and their consequences in many ocean regions from the top mixed layer to the abyss, information about flux Richardson number from controlled laboratory measurements is vital. The buoyancy flux ratio rises from 0 at small Richardson number to values of about 0.1 at Richardson number equal to 1, and then falls off for greater values. At Richardson number greater than 1, a wide range of power laws spanning the range from 0.5 to 1.5 are found to fit data for different experiments. The layering in this range causes flux clustering and may contribute to the relatively large scatter in the data and to the wide range of proposed power laws.
See also and time can influence a value. Thus, a maximum of RfC0.1 with a range of roughly 750% has remained unchanged over the last 20 years. However, so far most of the experiments have been conducted with Reynolds numbers of a few thousand or less, and with salt rather than temperature. Therefore the coverage in Re and Sc space is still limited. Finally there is evidence that a large Reynolds number might produce smaller maximum values (Figure 8), but comprehensive results are not yet available. A few studies have been conducted with two ingredients such as salt and temperature contributing density. If the value of Rf for each component differs, the phenomenon called ‘differential mixing’ is said to exist. At present, laboratory experiments do provide a small amount of evidence for differential mixing. Differential mixing is expected to be most important in oceanic regions where both temperature and salinity variations influence density, for instance, in polar regions. Therefore, for climate studies it would be important to incorporate differential mixing accurately in numerical models, especially since the salinity field is known to influence deep wintertime convection.
Differential Diffusion. Energetics of Ocean Mixing. Estimates of Mixing. Vortical Modes.
Further Reading Breidenthal RE (1992) Entrainment at thin stratified interfaces: The effects of Schmidt, Richardson and Reynolds numbers. Physics of Fluids A 10: 2141--2144. Cenedese C, Whitehead JA, Ascarelli TA, and Ohiwa M (2004) A dense current flowing down a sloping bottom in a rotating fluid. Journal of Physical Oceanography 34: 188--203. Fernando HJS (1991) Turbulent mixing in stratified fluids. Annual Review of Fluid Mechanics 23: 455--493. Park YG, Whitehead JA, and Gnanadesikan A (1994) Turbulent mixing in stratified fluids: Layer formation and energetics. Journal of Fluid Mechanics 279: 279--312. Rehmann CR and Koseff JR (2004) Mean potential energy change in stratified grid turbulence. Dynamics of Atmospheres and Ocean 37: 271--294. Turner JS (1973) Buoyant convection from isolated sources. In: Buoyancy Effects in Fluids, pp. 165--206. Cambridge, UK: Cambridge University Press. Turner JS (1986) Turbulent entrainment: The development of the entrainment assumption and its application to geophysical flows. Journal of Fluid Mechanics 170: 431--471.
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LAGOONS R. S. K. Barnes, University of Cambridge, Cambridge, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1427–1438, & 2001, Elsevier Ltd.
Introduction A ‘lagoon’ is any shallow body of water that is semiisolated from a larger one by some form of natural linear barrier. In ocean science, it is used to denote two rather different types of environment: ‘coastal lagoons’ (the main subject of this article) and lagoons impounded by coral reefs. Coastal lagoons are a product of rising sea levels and hence are geologically transient. Whilst they exist, they are highly productive environments that support abundant crustaceans and mollusks, and their fish and bird predators. Harvests from lagoonal aquaculture of up to 2500 kg ha1y1 of penaeid shrimp, of 400 tonnes ha1y1 of bivalve molluscs, and in several cases of 4200 kg ha1 y1 of fish are taken.
What is a Lagoon? Coastal lagoons are bodies of water that are partially isolated from an adjacent sea by a sedimentary barrier, but which nevertheless receive an influx of water from that sea. Some 13% of the world’s coastline is faced by sedimentary barriers, with only Canada, the western coast of South America, the China Sea coast from Korea to south-east Asia, and the Scandinavian peninsula lacking them and therefore also being without significant lagoons (Table 1, Figure 1). The lagoons behind such barrier coastlines range in size from small ponds of o1 ha through to large bays exceeding 10 000 km2. The median size has been suggested to be about 8000 ha. Although several do bear the word ‘lagoon’ in their name, many do not. Usage of the same titles as for fresh water habitats is widespread (e.g. E´tang de Vaccare`s, France; SwanPool, UK; Oyster Pond, USA; Lake Menzalah, Egypt; Benacre Broad, UK; Ozero Sasyk, Ukraine; Kiziltashskiy Liman, Russia, etc.), as is those for coastal marine regions (Peel-Harvey Estuary, Australia; Great South Bay, USA; Ringkbing Fjord, Denmark; Zaliv Chayvo, Russia; Pamlico Sound, USA; Charlotte Harbor, USA; and even GniloyeMore, Ukraine; Mer des Bibans, Tunisia; Mar Menor, Spain, etc.), whilst lagoons liable to hypersalinity are
often termed Sebkhas in the Arabic-speaking world (e.g. Sebkha el Melah, Tunisia). Conversely, a few systems with ‘lagoon’ in their name fall out with the definition; the Knysna Lagoon in South Africa, for example, is an estuarine mouth dilated behind rocky headlands between which only a narrow channel occurs. Coastal lagoons are most characteristic of regions with a tidal range of o2 m, since large tidal ranges (those 44 m) generate powerful water movements usually capable of breaching if not destroying incipient sedimentary barriers. Furthermore, the usual meaning of ‘lagoon’ requires the permanent presence of at least some water and large tidal ranges are likely to result in ebb of water from open systems during periods of low tide in the adjacent sea. Thus in Europe, for example, lagoons are abundant only around the shores of the microtidal Baltic, Mediterranean, and Black Seas. However, they are also present along some macrotidal coasts – such as the Atlantic north of about 471N (in the east) and 401N (in the west) (Figure 2) – where off shore deposits of pebbles or cobbles (‘shingle’) are to be found as a result of past glacial action. Here, shingle can replace the more characteristic sand of microtidal seas as the barrier material, as indeed it partially does in the lagoon-rich microtidal East Siberian, Chukchi, and Beaufort Seas in the Arctic, because it is less easily redistributed by tidal water movements. Nevertheless, sandy sedimentary barriers have developed (and persisted) in a few relatively macrotidal areas (e.g. in southern Iceland and Portugal), although the regions impounded to landwards retain water only during high tide for the reasons outlined above. Those environments that are true lagoons only during high tide are often referred to as ‘tidal-flat lagoons’. They are therefore the relatively rare macrotidal coast equivalent of the typical lagoons of microtidal seas. In many cases, the salinity of a lagoonal water mass is exactly the same as that of the adjacent sea, although where fresh water discharges into a lagoon its water may be brackish and a (usually relatively stable) salinity gradient can occur between rivermouth and lagoonal entrance channel. In regions where evaporation exceeds precipitation for all or part of a year, lagoons are often hypersaline. The second environment to which the word lagoon is applied is associated with coral reefs; coral here replaces the unconsolidated sedimentary barrier of the coastal lagoon. Circular atoll reefs enclose the ‘lagoon’ within their perimeter, whilst barrier reefs
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are separated from the mainland by an equivalent although less isolated body of water. Indeed barrier reef lagoons are virtually sheltered stretches of coastal sea. Atoll lagoons, however, are distinctive in being floored by coral sand which supports submerged beds of seagrasses and fringing mangrove swamps justas do the coastal lagoons of similar latitudes. The similarity between the two types of lagoon is nevertheless purely physiographic since the atoll lagoon fauna is that typical of coral reefs in general and not related in any way to those of coastal lagoons. Coral lagoons are covered in greater detail in the article Coral Reefs.
Table 1 The contribution of different continents to the world total barrier/lagoonal coastline of 3200 kma Continent
Percent of coastline barrier/lagoonal
Percent of world’s lagoonal resource
N. America Asia Africa S. America Europe Australia
17.6 13.8 17.9 12.2 5.3 11.4
33.6 22.2 18.7 10.3 8.4 6.8
a (Reproduced with permission from estimates by Cromwell JE (1971) Barrier Coast Distribution: A World-wide Survey, p. 50. Abstracts Volume, 2nd National Coastal Shallow Water Research Conference, Baton Rouge, Louisiana.)
The Formation of Lagoons Lagoons have existence only by virtue of the barriers that enclose them. They are therefore characteristic only of periods of rising sea level, when wave action can move sediment on and along shore, and – for a limited period of time – of constant sea level. At times of marine regression, they either drain or their basins may fill with fresh water. Many of the enclosing barriers have today been greatly augmented by wind-blown sand; for example, most of the larger systems along the South African coast (Wilderness, St Lucia, Kosi, etc.), and such enclosed lagoons have a particularly lake-like appearance (Figure 3). The precise physiographic nature of a lagoon then depends on the relationship of the barrier to the adjacent coastline. Starting at a point at which the barriers are some distance offshore, we can erect a sequence of situations in which the barriers are moved ever shorewards. This starting point can be exemplified by Pamlico and Albemarle Sounds in North Carolina, USA (Figure 4). There a chain of long narrow barrier islands located some 30 km off the coast enclosesan area of shallow bay of around 5000 km2. Further landwards movement of the barriers, often to such an extent that some of the larger barriers become attached to the mainland at one end to produce spits, leads to the situation currently characterizing most lagoonal coastlines and to the typical
Lagoonal coasts Figure 1 Major barrier/lagoonal coastlines.
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´ Etang de Kergalan
´ Etang de Trunvel
1 km Figure 2 Shallow lagoons along the coast of the Baie d’Audierne (Brittany, France) formed behind a shingle barrier beach. Sea water enters these systems only via overtopping of the barrier; lagoonal water leaves by percolation through the barrier.
coastal lagoon. Sea water then enters and leaves these lagoons through the channels between the islands or around the end of the spit. More rarely, the lagoons are enclosed by multiple tombolos connecting an offshore island to the mainland, or even tombolos joining a number of islands (as, for example, in Te Whanga, New Zealand). With the exception of tombolo lagoons, the long axis of typical lagoons – and many are extremely elongate – lies parallel to the coastline. Typical coastal lagoons are especially characteristic of microtidal seas. They range in size up to the 10 000 km2 Lagoa dos Patos, Brazil. The abundant lagoons that occur within river deltas form a special case of this category. Several deltas (e.g. those of the Danube and Nile) have developed within – and have now largely or completely obliterated – former lagoons enclosed by spits and intervening barrier island chains. The filling of lagoons with sediment when they receive river discharge is an inevitable process and many surviving examples are but small remnants of their past extents (Figure 5) – islands of water in a sea of marshland.
Some 5000 years BP, the Lake St Lucia lagoon in South Africa, for example, had a length in excess of 110 km and a surface area of nearly 1200 km2; infilling with sediment has reduced the body of free water to a length of 40 km and an area of 310 km2 (see Figure 3). Allowing for the shallowing that has also taken place, this is equivalent to an average input of 623 000 m3 of sediment per year. Infill, coupled with barrier transgression via wash-over fans, is probably the ultimate fate of virtually all of the world’s lagoons. The next stage in the hypothetical sequence sees the barrier very close to the mainland, if not plastered onto it. This has several consequences, of which the first is impingement on emerging river systems. ‘Estuarine lagoons’, which as their name implies merge in to estuaries, have formed where barriers have partially blocked existing drowned river valleys. For this reason they usually have their long axis perpendicular to the coastline. Even if blockage of the river is complete, lagoonal status may remain if sea water can enter by overtopping of the barrier during high water of spring tides.
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Enkovugeni (8_35) Makhawulani (4_22) Mpungwini (1_ 13) 10 km
Nhlange (1_ 6)
Amanzimnyama (fresh)
Dunes
Marsh
Figure 3 The Kosi Lakes (KwaZulu/Natal, South Africa), with approximate salinity ranges (in practical salinity units) in parentheses. They are enclosed within Pleistocene sand dune fields, and form a classic case of ‘segmentation’ of an original linear lagoon into aseries of rounded basins. (Salinity data reproduced with permission from Begg G (1978) The Estuaries of Natal. Natal Town and RegionalPlanning Report.).
Nevertheless, many former estuarine lagoons are now completely fresh water habitats with seawater entry being prevented by the barrier. In many tropical and subtropical regions, closure of the estuarine lagoons is seasonal. During the wet season, river flow is sufficient to maintain breaches through the barriers. In the dry season, however, flow is reduced and drift can seal the barrier. If fresh water still flows, the whole isolated basin then becomes fresh for several months. The mouths of many natural lagoonal barriers are periodically breached by man for a variety of reasons ranging from ensuring the entry of juveniles of commercially important mollusks, crustaceans, and fish to temporarily lowering water levels to avoid damage to property. Longshore movement of barrier sediment may also divert the mouths of discharging estuaries several kilometers along the coast, as for example the 30 km deflection of that of the Senegal River by the Langue de Barbarie in West Africa, and the creation of the
Indian River Lagoon in Florida, USA. Not infrequently a new mouth is broken through the barrier, naturally or more usually as a result of human intervention, and the diverted stretch can become a dead-end backwater lagoon fed by backflow. Other consequences of on shore barrier migration also involve human intervention. Many small scale lagoons have (unwittingly) been created by land reclamation schemes. Regions to landwards of longshore ridges that were once, for example, lowlying salt marsh but which are now reclaimed usually receive an influx of water from out of the barrier water-table and this collects in depressions that were once part of the creeks draining the marshes. Equivalently, gravel pits or borrow pits in coastal shingle masses, from which building materials have been extracted, similarly receive an influx of water from out of the shingle. In both cases, the shingle water-table is derived from sea water soaking into the barrier during high tide in the adjacent sea
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Albemarle Sound
Pamlico Sound
Cape Hatteras
Cape Hookout
100 km Figure 4 Pamlico and Albemarle Sounds (North Carolina, USA).
together with such rainfall as has also soaked in. Some of the natural limans of the Black Seacoast also receive their salt input via equivalent seepage. Finally in this category, barriers may come onshore in such a fashion as to straddle the mouth of a preexisting bay. ‘Bahira lagoons’ (from the Arabic for ‘littlesea’) are pre-existing partially land-locked coastal embayments, drowned by the post glacial rise in sea level, that have later had their mouths almost completely blocked by the development of sedimentary barriers. Also included here are systems in which the sea has broken through a pre-existing sedimentary barrier to flood part of the hinterland, but inwhich the entrance/exit channel remains narrow. Bahira lagoons clearly merge into semi-isolated marine bays. Although lagoons may come and go during geological time, some of the larger lagoon systems are not solely features of the present interglacial period. Some, including the Lagoa dos Patos, the Gippsland Lakes of Victoria, Australia, and the Lake St Lucia lagoon, may incorporate elements of barriers and basins formed during previous marine transgressions.
The history of the Gippsland lakes over the last 70 000 years has been reconstructed in some detail (Figure 6). Fossil assemblages of foraminiferans suggest that lagoons may have been a feature of the Gulf Coast of the USA and Mexico at intervals ever since the Jurassic.
Lagoonal Environments It can be argued that the main force structuring lagoons is the extent to which they are connected to the adjacent sea, and they have been divided into three or four types on this basis(Figure 7) which largely reflect points along the hypothetical evolutionary sequence described above. ‘Leaky lagoons’ function virtually as sheltered marine bays (see also Figure 4). As a result of large tidal ranges in the adjacent sea or the existence of many and/or wide connecting channels, interactions with the ocean are dominant and the lagoons have tidally fluctuating water levels, short flushing times and a salinity equal to that ofthe local sea water.
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Former lagoon
Present delta
50 km
Figure 5 Infill of a former lagoon (established around 6000 years ago) on the Romanian/Ukrainian coast of the Black Sea with sediment carried by the River Danube. The current Danube delta occupies virtually all the former lagoon.
‘Restricted lagoons’ have asmall number of narrow entrance channels through elements of a barrier is land chain or enclosing spits and are therefore more isolated from the ocean(see also Figure 7). Characteristics include: longer flushing times, a vertically well-mixed water column, both wind and tidal water movements as forcing agents, and brackish to oceanic salinity. ‘Choked lagoons’ are connected to the ocean by a single, often long and narrow, entrance channel that serves as a filter largely eliminating tidal currents and fluctuations in water level(see also Figure 4). The 1000 km2 Chilka Lake lagoon, for example, communicates with the Bay of Bengal via a channel 8 km long and 130 m across at its widest. Choked lagoons are typical of coasts with high wave energy and significant longshore drift. Characteristics include: very long flushing times, intermittent stratification of the water column by thermoclines and/or haloclines, wind action as thed ominant forcing agent, purely freshwater zones near inflowing rivers, and otherwise brackish, and in some climatic zones periodic hypersaline, waters. Wet season rainfall in the 44 km2 Laguna Unare, Venezuela (at the end of the dry season), for example, can change the salinity from 60–92 to 18–25, at the same time increasing lagoonal depth by 1 m, area by 20 km2 and volume by 76 106m3.
The ‘closed lagoons’ shown in Figure 2 form the limiting condition of effective isolation from direct inputs from the ocean except via occasional overtopping or after percolation through the barrier system. Most are not only ‘former lagoons’, but also current coastal freshwater lakes; for example, the much studied Slapton Ley in England. Characteristically restricted and choked lagoons are very shallow (usually about 1 m and almost always o5 m deep), floored by soft sediments, fringed by reedbeds(Phragmites and/or Scirpus), mangroves (especially Rhizophora) or salt marsh vegetation, and support dense beds of submerged macrophytes such as seagrasses (e.g. Zostera), pondweeds (Potamogeton) or Ruppia, together with green algae such as Chaetomorpha. The action of the dense beds of submerged vegetation may be to raise levels of pH and to contribute considerable quantities of organic matter. In stratified choked lagoons, decomposition of this organic load below the thermocline or halocline can then lead to anoxic conditions, not withstanding surface waters super saturated with oxygen (Figure 8).
Lagoonal Biotas and Ecology Insofar as is known, although the species may be different, the ecology of coastal lagoons shows the
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Lag
oon
Prior
barrie
r
(A)
Inner barrier (B)
(C)
(D)
Outer barrier 25 km
(E) Pleistocene land surface
Alluvium and swamps
Outlet Barriers
Figure 6 Evolution of the Gippsland Lagoon system (Victoria, Australia). (Reproduced with permission from Barnes 1980; after Bird ECF (1966) The evolution of sandy barrier formations on the East Gippsland coast. Proceedings of the Royal Society of Victoria 79:75–88.) The prior barrier (A) can be dated to about 70 000 years BP; the inner barrier (B) was formed during the late Pleistocene; the situation in (C) represents a low sea level phase of the late glacial, and (D) and (E) further evolution during the current marine transgression.
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latitude lagoons of comparable nature; and whilst shallow phytoplankton-dominated lagoons are more productive than deeper ones (depth 42 m), the converse may be true of macrophyte-dominated systems. An order of magnitude calculation (assuming that the total area of the world’s lagoons is around 320 000 km2 and that average lagoonal productivity is 300 g Cm2 y1) yields a total annual lagoonal production of 1011 kg fixed C. The contribution of lagoons to total oceanic carbon fixation is therefore minor, although they are as productive per unit area as are estuaries, and are only less productive than some regions of upwelling, coral reefs, and kelp forests. What happens to this primary production? Data are still scarce. There has been much argument as to whether estuaries are sources or sinks for fixed
10
0
km
same general pattern as seen in estuaries and other regions of coastal soft sediment, although the relative importance of submerged macrophytes may be greater in lagoons (Figure 9). The action of predators as important forces structuring the communities in lagoons as in similar areas, for example, is reflected by the young stages of many species occurring within the dense beds of submerged macrophytes, as the hunting success of predatory species is lower there. The primary productivities of lagoons appear to vary with their general nature, their latitude, and their depth. Choked lagoons tend to be the most productive, with recorded maxima approaching 2000 g C m2 y1, and with productivities (allowing for the effect of latitude) some 50% more than inrestricted lagoons. Temperate zone lagoons achieve only about 50–70% of the productivity of low
15 km
Choked
Restricted
Leaky
25 km Figure 7 ‘Choked’ (Lagoa dos Patos, Brazil), ‘restricted’ (Laguna di Venezia, Italy) and ‘leaky’ (Laguna Jiquilisco, El Salvador) lagoons.
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385
0
1 Depth (m)
Halocline
Oxycline
2
3 0
5
10
0
15
Salinity
100
50 Oxygen (% saturation)
Figure 8 A halocline and consequent oxycline in the choked Swanpool Lagoon, England, after a prolonged period of high freshwater input (salinity inpractical salinity units). (Reproduced with permission from data in Dorey AE et al. (1973) An ecological study of the Swanpool, Falmouth II. Hydrography and its relation to animal distributions. Estuarine Coastla and Marine Science 1: 153–176.)
Pelican, osprey, fish-eagle, etc.
Mullet
Flamingo
Shore-birds, heron, etc.
Large fish
Waterfowl
Crabs
Amphipod crustaceans, gastropod mollusks, etc.
Prawns
Polychaete worms
Small fish
Bivalve mollusks
Algal productivity and detrital materials Macrophyte material Figure 9 Simplified lagoonal food web.
carbon. In so far as information is available it appears that lagoons are more nearly in balance, as is appropriate for their more isolated nature, although all those so far studied do function as slight sinks.
Even though relatively isolated, however, no lagoon is self-contained, not least in that many elements in the faunas migrate between lagoons and other habitats.
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Coastal lagoons are heavily used by wetland and shore birds as feeding areas, most noticeably by grebes, pelicans, ibis, egrets, herons, spoonbills, avocets, stilts, storks, flamingoes, cormorants, kingfishers, and fish eagles. Such birds prey on the abundant lagoonal invertebrate and fish faunas. These comprise mixtures of three different elements: both (a) essentially freshwater and (b) essentially marine and/or estuarine species capable of with standing adegree of brackishness, as well as (c) specialist lagoonal or ‘paralic’ species that are not in fact restricted to lagoons but also (or closely related species) occur in habitats like the Eurasian inland seas (e.g. the Caspian and Aral), as well as in (the usually man-made) tideless brackish ponds and drainage ditches that are abundant in reclaimed coastal regions. The main feature of the lagoonal species category is that they are species of marine ancestry that seem to be restricted to shallow, relatively tideless maritime habitats; over their ranges as a whole they are also clearly capable of inhabiting a wide range of salinity, including full-strength and concentrated sea water, but not fresh water. Their relatedness to marine species is evidenced by the fact that they often form species-pairs with marine or estuarine species. The freshwater component of lagoonal faunas is often greater than would be expected in comparable salinities in estuaries, with the common occurrence of adult water beetles and hemipteran bugs besides the omnipresent dipteran larvae. The reasons for this are not understood, but may relate to decreased water movements. Because oceanic and freshwater inputs into alagoon occur at different points around its perimeter, these three elements of the fauna tend to be broadly zoned in relation to these inputs. Indeed, an influential French school, the ‘Group d’Etude du Domaine Paralique’, identifies six characteristic biotic zones in lagoons based not upon salinity but upon ‘a complex and abstract value which cannot be measured in the present state of knowledge’ which the group terms ‘confinement’. This is a function, at least, of the extent to which the water mass at any given point is isolated from oceanic influence. Beginning with the Mediterranean Sea coast, disciples of this school have now published maps of the location of the ‘six degrees of confinement’ in many lagoons throughout the world. Lagoons also form the nursery grounds and adult feeding areas for a large number of commercially important fish and crustaceans that migrate between this habitat and the sea, most of which spawn outside the lagoons and only later move in actively or in some cases probably passively. In respect of the fish, temperate regions are amongst others used byeels
(Anguillidae), sea bass (Moronidae), drum (Sciaenidae), sea bream (Sparidae), grey mullet (Mugilidae), flounder (Pleuronectidae), and various clupeids, and in warmer waters these are joined by many others, including cichlids, milkfish (Chanidae), silver gars (Belonidae), grouper (Serranidae), puffer fish (Tetraodontidae), grunts(Pomadasyidae), rabbit fish (Siganidae), various flatfish (Soleidae, Cynoglossidae), and rays (Dasyatidae). Therefore, throughout the world lagoons support (often artisanal) fisheries, as well as being the location of bivalve shellfish culture (mussels, oysters, and clams). For obvious reasons, most data on secondary productivity have been collected in relation to these fisheries. The fish yield from lagoons is greater than from all other similar aquatic systems, averaging (n ¼ 107) 113 kg ha1 y1 and with a median value of 51 kg (Table 2) and with 13% exceeding 200 kg ha1 y1. Yields vary with the intensity of human involvement in the catching and stocking processes. In the Mediterranean Sea, the region for which most data are available, the yield without intensive intervention is some 82 kg ha1y1; with the installation of permanent fish traps this increases to 185 kg; and with the addition of artificial stocking with juvenile fish it increases to 377 kg. For a 10 year period, the fishermen’s cooperative based on the Stagno di Santa Giusta in Sardinia harvested nearly 700 kg ha1 y1. West African lagoons with ‘fish parks’ – areas that attract fish because of the provision of artificial refuges – are considerably more productive, with average fish yields of 775 kg ha1 y1. Amongst the invertebrates, penaeid shrimp, mudcrabs (Scylla), oysters, mussels, and the arcid ‘cockle’ Anadara are the major harvested organisms. Yields of Penaeus may attain 2500 kg ha1 y1in lagoonal systems of aquaculture, whilst those of the oyster Crassostrea and the mussel Perna can both be 400 tonnesha1 y1. In the Indian subcontinent wholecommunities can be economically entirely dependent
Table 2 Mean fisheries yield from coastal lagoons in relation to those from other aquatic habitatsa Habitat
Fish yield (kg ha1y1)
Percent of sites exceeding yield of 200 kg ha 1y 1
Coastal lagoons Continental shelf Coral reef Fresh waters
113 59 49 34
13 5 0 2
a
(Data reproduced with permission from Chauvet, 1988.)
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Table 3 Discharges in the early 1990s into two lagoons on the southern (Polish/Russian/ Lithuanian) coast of the Baltic Sea(inputs in tonnes yr1)a Parameter
Curonian Lagoon
Vistula Lagoon
Surface area Mean depth Maximum depth Volume Annual water exchange BOD input Inorganic N Inorganic P Chlorinated substances Cu Zn Pb
1584 km2 3.8 m 5.8 m 6.0 km3 27.3 km3 150 000 20 000 5000 1 30 50 600
838 km2 2.6 m 5.2 m 2.3 km3 ? 47 000 25 000 2100 ‘High concentrations’ ? ? ?
a (Reproduced with permission from data in Coastal Lagoons and Wetlands in the Baltic. WWF Baltic Bulletin 1994, no 1.)
on these shrimp, crab, and fish harvests; for example, the 60 000 people living around Chilka Lake, India. However, many of the world’s lagoons are threatened habitats, suffering deterioration as a result of pollution and destruction, through reclamation and from the natural losses resulting from succession to freshwater and swamp habitats and from landwards barrier migration. The discharge of materials into lagoons and its consequences are essentially similar to those in any other semi-enclosed coastal embayment, but there is one specifically lagoonal feature. As they are often used for aquaculture and yields are generally related to primary productivity, lagoonal waters are frequently deliberately enriched to boost catches. Such enrichment varies from domestic organic wastes from the surrounding communities to commercial processed fish foods. Probably in the majority of such cases, however, the result of this nutrient injection has been eutrophication, loss of macrophytes, deoxygenation, and in several areas a change in the primary producers in the direction of a bacteria-dominated plankton and, across wide areas, benthos as well. In Mediterranean France, this all too frequent state of affairs is known as ‘malai¨gue’. Culture of mussels in the Thau Lagoon in France produces an input to the benthos of some 45 000 tonnes (dry weight) of pseudofecal material. Not surprisingly, at times of minimum throughput of water, malai¨gues can cause mass mortality of the cultured animals and degradation of the whole habitat. Thus in Europe, malai¨gues in the south, pollution in the Baltic lagoons(Table 3), and reclamation ofthose on the Atlantic seaboard have rendered the habitat especially threatened even at a continental level. For this reason they are now a ‘priority habitat’ under the European Union’s
Habitats Directive. Intensiv elagoonal aquaculture also injects not only nutrients, but antibiotics, hormones, vitamins, and a variety of other compounds, and the wider effects of these are giving cause for concern.
See also Crustacean Fisheries. Fish: Demersal Fish (Life Histories, Behavior, Adaptations). Eels. Eutrophication. Geomorphology. Macrobenthos. Mangroves. Molluskan Fisheries. Pelecaniformes. Phytobenthos. Primary Production Distribution. Salt Marshes and Mud Flats.
Further Reading Ayala Castan˜ares and Phleger F.B. (eds.) (1969) Lagunas Costeras. Un Simposio. Universidad Nacional Auto´noma de Me´xico. Barnes RSK (1980) Coastal Lagoons. Cambridge: Cambridge University Press. Bird ECF (1984) Coasts, 3rd edn. London: Blackwell. Chauvet C (1988) Manuel sur l’Ame´nagement des Peˆches dans les Lagunes Coˆtie`res: La Bordigue Me´diterrane´enne. FAO Document Technique sur les Pches No. 290. FAO: Cooper JAG (1994) Lagoons and microtidal coasts. In: Carter RWG and Woodroffe CD (eds.) Coastal Evolution, pp. 219--265. Cambridge: Cambridge University Press. Emery KO (1969) A Coastal Pond Studied by Oceanographic Methods. Elsevier. Gue´lorget O, Frisoni GF, and Perthuisot J-P (1983) La zonation biologique des milieux lagunaires: de´finition d’une e´chelle de confinement dans le domaine paralique
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mediterrane´en. Journal de Recherche Oce´anographique, Paris 8>: 15--36. Kapetsky JM and Lasserre G (eds.) (1984) Management of Coastal Lagoon Fisheries, vol. 2 vols.. Rome: FAO. Kjerfve B (ed.) (1994) Coastal Lagoon Processes. Elsevier. Lasserre P and Postma H (eds) (1982) Les Lagunes Ctie`res. Oceanologica Acta (Volume Spe´cial.) Rosecchi E and Charpentier B (1995) Aquaculture in Lagoon and Marine Environments. Tour du Valat.
Sorensen J, Gable F, and Bandarin F (1993) The Management of Coastal Lagoons and Enclosed Bays. American Society of Civil Engineers. Special issue (1992) Vie et Milieu 42(2): 59--251. Ya´n˜ez-Arancibia A (ed.) (1985) Fish Community Ecology in Estuaries and Coastal Lagoons. Universidad Nacional Auto´noma de Me´xico.
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LAGRANGIAN BIOLOGICAL MODELS D. B. Olson, C. Paris, and R. Cowen, University of Miami, Miami, FL, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1438–1443, & 2001, Elsevier Ltd.
Introduction The Swiss mathematician Leonhard Euler (1707– 1783) derived the formulations for describing fluid motion by either measuring the properties of the fluid at a fixed point overtime or alternatively following the trajectory of a parcel of fluid as it is carried with the flow. The first of these is known as the Eulerian description of the flow, while the method following a material parcel or particle is known as the Lagrangian description after the French mathematician Joseph Lagrange (1736–1813). Most of the theory used to model ocean currents is posed in an Eulerian frame because of the difficulties in solving the momentum equations in the complicated matrices that arise in the Lagrangian form of the equations. However, it is often useful to use the Lagrangian frame of reference when considering the manner in which mixing occurs in turbulent flows such as those found in the oceans. It is also common to measure these flows by using drifters or floats that trace out oceanic currents. As shown below, the Lagrangian description is also conducive to handling models of marine populations in many cases. This is especially true when the models include quantities that structure the population, such as age, genetics, or physiological traits that depend upon the history of individual organisms that are carried in or swim through oceanic flows. Organisms that drift freely with the currents are termed planktonic, while those that can swim effectively are termed nektonic fauna. Here Lagrangian methods for considering populations of both plankton and nekton are given. Much of the detailed formalism can be found in Okubo (1980) (see Further Reading). The present discussion highlights the application of these methods to marine population dynamics.
Comparing the Eularian and Lagrangian Formulations To understand the difference between the Lagrangian and Eulerian formulations, consider the population dynamic equations for marine organisms. If the ith
population is made up of Ni individuals, one can write an equation for each individual. This will include each organism’s position, Xm ðtÞ, as a function of time t. The total, or Lagrangian, derivative of Xm ðtÞ with respect to time, dXm/dt, gives the individual’s velocity, Vm(t). This can be separated into the influence of the advection of the organism by ocean currents, U (Xm, t), and a swimming contribution, Us(Xm, B,t), where B (Xm, t) is a vector of behavioral clues. These clues involve both physical and biotic components of the environment. The acceleration of the individual organism is then derived by carrying out another differentiation in time (eqn [1]). dVm @U @Us ¼ þ UdrVm þ Us drVm ¼ FðBÞ þ dt @t @t
½1
r is the spatial gradient operator and F is the gravitational and viscous forces imposed on the organism as well as behavioral responses, i.e., swimming. Notice that the Lagrangian derivative on the left leads to a set of Eulerian terms that involve spatial gradients in the fluid velocity and the behavioral clues on the right side of the equation. This equation fully expresses the motion of an individual. To the equation describing the organism’s motion, a set of state relations must be added expressing changes in its physiological state, its age or stage, and the probabilities of its death and reproduction to explain population dynamics. Such a model considering the conditions of each individual explicitly in a population is called an individual-based model (IBM). Individual-based models provide a method for understanding behavior and small groups of organisms as discussed below. For large populations, however, the number of equations involved becomes impossible to handle. It is therefore common to introduce the concept of organism density, ni ¼ Ni/A, where Ni is the number of individuals and A is a given a real measure. The density of the ith taxon is then measured in numbers per square kilometer of ocean surface area. This leads to a continuous spatial field equation. It is typical to consider the mean field of population density and perturbations about it so ni ¼ /ni S þ n0i . Here the first term is the mean population density and the second the perturbations (or variance) about the mean; the mean is over the population. The same separation can be done for the velocity components such that U ¼ /US þ U0 and
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LAGRANGIAN BIOLOGICAL MODELS
Day 10 22°N
20°N
20°N
18°N
18°N
Longitude
Longitude
Day 1 22°N
16°N
16°N
14°N
14°N
12°N
12°N
10°N
75°W
70°W 65°W Latitude
10°N
60°W
75°W
60°W
70°W 65°W Latitude Day 30
Day 20 22°N
22°N
20°N
20°N
18°N
18°N
8
Longitude
Longitude
6
16°N
16°N
14°N
14°N
12°N
12°N
4
2
10°N (A)
75°W
70°W 65°W Latitude
60°W
10°N
75°W
65°W 70°W Latitude
60°W
0 log10 ni
22°N
20°N
Latitude
18°N
16°N
14°N
12°N
10°N 75°W (B)
70°W
65°W Longitude
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60°W
55°W
LAGRANGIAN BIOLOGICAL MODELS
Us ¼ /Us S þ Us 0 . Equation [1] above involves products of velocity components with each other such that the contributions to the mean motion of the population and therefore its average spread will involve both the mean velocities and correlations between velocity perturbations. In the field equations these correlation’s between fluctuations in the turbulent fluid velocity or swimming behavior will lead to turbulent and behavioral diffusion. It is typical to introduce diffusivities, K for the turbulence and Ks for the behavioral related dispersion. There is also a correlation between the variations in the environmental factors, including the distance to members of the same species that come into effect given B ¼ /BS þ B0 . The resulting field equation for the expected mean density of an organism is then given by eqn [2]. dni @ni ¼ þ Vm drni ¼ FðBÞ þ r½ðk þ ks Þrni dt @t
½2
The k are inside the spatial gradient operators to denote that they are functions of space. Taking the ks term all the way inside the r operators allows density or schooling effects on population density through behavioral preferences for nearest neighbor distance. The results of the field model versus the Lagrangian model following individuals can lead to impressive differences (Figure 1). The diffusion model in Figure 1A using the equations above lead to a finite probability of finding organisms everywhere in a domain immediately. The Lagrangian treatment limits an organism’s spread to the fastest velocities present, so that it takes a finite time for spread. It is important to note, however, that there must also be population losses that are more abrupt than simple exponential decrease of the population across habitat boundaries. This occurs in most population parametrizations such as a logistic growth with linear mortality to achieve realistic population distributions. In the real ocean, while it takes much longer than in the analytical diffusion model, the Lagrangian motions will still lead to a finite possibility of finding organisms everywhere in the domain at large timescales, unless mortality is properly treated.
391
Simple Models in the Lagrangian Frame The most important issue in modeling marine populations is providing an accurate depiction of the physics, the biology, and the intricate biological/ physical interactions that occur. These are represented in the equations above by mean quantities acting on means or perturbations and then by correlations between both physical and biological perturbations. One use of the Lagrangian description of motion is to simulate these interactions along a fluid trajectory in the case of the plankton (Figures 2 and 3). In this case a simple meandering current and its impact on populations is envisioned. The calculation involved is a simple integration that in Figure 2 reveals the basic response without biological non linearities. In Figure 3 the impact of a primary production response as seen in Figure 2 on a density-dependent population conceals a set of more interesting patterns, including extinction. The situation in Figures 2 and 3 involves dynamics that allow exact calculations, i.e, in this case simple integrations of the functions without any use of numerics. This sort of analysis is recommended for testing morecomplicated cases where numerical methods become a major issue. Essentially these simple applications use the Lagrangian frame of viewing advection as a means of allowing simple calculations of population dynamics. The Lagrangian frame becomes indispensable when structured populations are considered.
Simulations of Populations with Demographic Structure The Lagrangian description of the path that biological entities follow through the ocean environment becomes the only feasible method for treating population dynamics where the detailed history of the populations’ interaction with the physical environment and other populations are important. Early works in this area include the plankton models of Wolf and Woods following the details of mixed layer and thermocline development in the North Atlantic over many seasons and the work on zooplanktonin the coastal environment by Hoffman et al. The problem becomes that the Lagrangian
Figure 1 (Left) (A) Simulation of larval reef fish drift from Barbados using the mean flow into the eastern Caribbean and a k of 5000 m2 s1 at 1, 10, 20, and 30 days after spawning. A typical larval mortality rate m ¼ 0.2 (or B18% per day) is applied to larval abundance (Ni). (B) Lagrangian simulation of the same case with trajectories computed from an oceanic general circulation model at day 30 (Cowen et al., 2000). The survivors are indicated by red dots after applying the same m ¼ 0.2. Note that none are on suitable island habitat after 30 days. For the diffusive case (A) there is a finite probability of finding larvae well beyond the range of any of the simulated trajectories. In this case the mortality is truncated by the 30-day duration of planktonic behavior in the fish’s assumed development, i.e., after this time there is assumed recruitment to an island or death.
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392
LAGRANGIAN BIOLOGICAL MODELS
1.4
Meander phase
K
1.3 1.2 1.1 1.0 0
1 2 (Lagrangian period)
3
Figure 2 Response of phytoplankton to a simple meandering current. The meander is assumed to set up a simple sinusoidal series of upwelling (high) and down welling (low) at meander crests and troughs, respectively, that provide nutrients upon the upwelling phase. The model assumes logistic phytoplankton response to a sinusoidal carrying capacity (K) and is solved analytically by simple integration of the Lagrangian equation in time. The forcing function involves linear response to a cosine shown in the figure.
frame (i.e., following individual trajectories of individuals or individual subpopulations) is the only way to track the biological dynamics where the past history plays a major role in determining the present dynamics. These historical parameters may involve the past history of nutrient or forage availability, the temperature and salinity encountered over the development of the organisms, or the past history of selection on the genetic structure of populations when reproduction occurs. As an example of a structured population simulation in a turbulent ocean gyre, a simulation of a population of physiologically and genetically structured pelagic copepods is described. The population
29 y 19 y 120 9y 100
Biomass
Phytoplankton biomass
1.5
80
60 40
K
2.5
20 0 0
80°
120° 160° 200° 240° 280° 320° 360°
Northern
K1
Southern
1.1530
1.1570
1.5
× 10
1.0
K2 K3
0.5
0
Growth rate
Phytoplankton biomass
40°
Southern
(A)
2.0
20
1.1560
10
1.1550
Flow
0
1 2 (Lagrangian period)
0
3
Figure 3 A calculation of the zooplankton response to the phytoplankton distribution in a meander like that treated in Figure 2. Here the integration in time along trajectories is slightly more complicated but still analytical. The zooplankton response is parametrized by a Hollings type 2 curve such that the time dependence of zooplankton (Z) is governed by the equation below.
dZ ZP ¼r dZ dt K0 þ P Here r is a growth rate, P is the sinusoidal phytoplankton field, K0 is a half-saturation term, and d is the death rate for Z. The pattern of the carrying capacity, K is shown at the top of figure. Four different Ks are shown with different magnitudes K1–K3 and fourth that goes extinct. See Olson and Hood (1994) and the literature cited there for further discussion of meander impacts on marine ecosystems.
0
(B)
40°
Southern
80°
1.1540 120° 160° 200° 240° 280° 320° 360°
Northern
Southern
Figure 4 (A) Biomass of copepods per m2 of surface area as a function of distance (y) around the gyre, at 10-year intervals. Gyre circulation time is 3 years. The carrying capacity at the northern and southern ends of the gyre are indicated by dashed lines. Note that mixing can cause a region to exceed the carrying capacity. The variation in populations is largest at the highest carrying capacity. (B) The mean growth rate (biomass per day) in different segments around the gyre at the three times. The scale at 9 years is 10 times that for the later times (scale at right). Growth rate is going slowly to a constant or fixed state as expected from population genetics grounds. The distribution of growth rates matches the cosine nature of the carrying capacity distribution, but is fully advective in the sense that these patterns advect with the mean flow around the gyre. The direction of this drift is indicated by an arrow.
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LAGRANGIAN BIOLOGICAL MODELS
is based on the properties of Nanocalanus minor, a copepod found in the subtropical gyre in the North Atlantic. The model is designed to consider the dynamics behind the mitochondrial DNA patterns found in this population in the Gulf Stream. The Gulf Stream and its recirculation gyre are treated as a circular flow with superimposed turbulence. The copepods are simulated as subpopulations, each carried on Lagrangian particles advected in this flow. The populations are subject to a carrying capacity, K ¼ K0 þ K0 sinðy=2Þ, that is high in the northern Gulf Stream and low in the oligotrophic southern portions of the gyre. The population has variable growth rates that are controlled genetically. The growth potential is determined by the statistics of the local breeding subpopulations and by selection induced by competition at a given location and time for food. Selection is local since there is not an optimal growth rate in the sense that it pays to have a high reproductive potential in the northern gyre under low population densities. The offspring of such a population are inevitably at a disadvantage, however, when advected into the southern oligotrophic reaches of the gyre. Since the resulting genetic and physiological attributes at a location depend upon the past history of all of the subpopulations contributing to the interaction at a given time, this sort of simulation becomes computationally impossible in a Eulerian frame. The population distributions in Figure 4, done in a Lagrangian simulation, takes only an hour on a laptop computer. The details of a suite of such simulations are currently being compared to population density and gene sequences.
Conclusions The use of Lagrangian particle-following simulations in modeling population dynamics allows several advantages over Eularian fixed-grid calculations. For simple models the advantage is that the population equations can be simply integrated in time. As new techniques for tracking fluid parcels and therefore planktonic trajectories or individual large pelagic fish or whales become more available, models using real trajectories will become possible. The other advantage that direct Lagrangian simulation of turbulent dispersal of organisms has is that it overcomes the problems that advection/diffusion schemes have with population densities at large distances from their source. Finally, the largest promise in Lagrangian
393
simulations is their use in models that explicitly treat the demographic traits of populations. With the everincreasing realism in physical models of the marine environment and Lagrangian population models, new insights into marine population dynamics are possible.
See also Plankton Viruses. Population Dynamics Models.
Further Reading Bucklin A, LaJeunesse TC, Curry E, Wallinga J, and Garrison K (1996) Molecular genetic diversity of the copepod, Nannocalanus minor: genetic evidence of species and population structure in the N. Atlantic Ocean. Journal of Marine Research 54: 285--310. Carlotti F (1996) A realistic physical-biological model for Calanus finmarchicus in the North Atlantic. A conceptual approach. Ophelia 44: 47--58. Carlotti F and Wolf KU (1998) A Lagrangian ensemble model of Calanus finmarchicus coupled with a 1-D ecosystem model. Fishers and Oceanography 7(7): 191--204. Cowen RK, Lwiza KMM, Sponaugle S, Paris CB, and Olson DB (2000) Connectivity of marine populations: Open or closed? Science 287: 857--859. Flierl G, Gru¨nbaum D, Levin S, and Olson DB (1999) From individuals to aggregation: the interplay between behavior and physics. Journal of Theoretical Biology 196: 397--454. Hoffmann EE, Halstrom KS, Moisan JR, Haidvogel DB, and Mackas DL (1991) Use of simulated drifter tracks to investigate general transport patterns and residence times in the coastal transition zone. Journal of Geophysical Research 96: 15041--15052. Metz JAJ and Diekmann O (eds.) (1986) The Dynamics of Physiologically Structured Populations, 68. Berlin: Springer-Verlag. Okubo A (1980) Diffusion and Ecological Problems: Mathematical Models, 10. Berlin: Springer-Verlag. Olson DB and Hood RR (1994) Modelling pelagic biogeography. Progress in Oceanography 34: 161--205. Wolf KU and Woods JD (1998) Lagrangian simulation of primary production in the physical environment: The deep chlorophyll maximum and nutricline. In: Rothschild BJ (ed.) Toward a Theory on Biological– Physical Interactions in the World Ocean, pp. 51--70. Dordrecht: Kluwer Academic.
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LAND–SEA GLOBAL TRANSFERS F. T. Mackenzie and L. M. Ver, University of Hawaii, Honolulu, HI, USA
and groundwater flows, thus moving toward the sea in a greater degree than before.
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The Changing Picture of Land–Sea Global Transfers Introduction The interface between the land and the sea is an important boundary connecting processes operating on land with those in the ocean. It is a site of rapid population growth, industrial and agricultural practices, and urban development. Large river drainage basins connect the vast interiors of continents with the coastal zone through river and groundwater discharges. The atmosphere is a medium of transport of substances from the land to the sea surface and from that surface back to the land. During the past several centuries, the activities of humankind have significantly modified the exchange of materials between the land and sea on a global scale – humans have become, along with natural processes, agents of environmental change. For example, because of the combustion of the fossil fuels coal, oil, and gas and changes in land-use practices including deforestation, shifting cultivation, and urbanization, the direction of net atmospheric transport of carbon (C) and sulfur (S) gases between the land and sea has been reversed. The global ocean prior to these human activities was a source of the gas carbon dioxide (CO2) to the atmosphere and hence to the continents. The ocean now absorbs more CO2 than it releases. In preindustrial time, the flux of reduced sulfur gases to the atmosphere and their subsequent oxidation and deposition on the continental surface exceeded the transport of oxidized sulfur to the ocean via the atmosphere. The situation is now reversed because of the emissions of sulfur to the atmosphere from the burning of fossil fuels and biomass on land. In addition, river and groundwater fluxes of the bioessential elements carbon, nitrogen (N), and phosphorus (P), and certain trace elements have increased because of human activities on land. For example, the increased global riverine (and atmospheric) transport of lead (Pb) corresponds to its increased industrial use. Also, recent changes in the concentration of lead in coastal sediments appear to be directly related to changes in the use of leaded gasoline in internal combustion engines. Synthetic substances manufactured by modern society, such as pesticides and pharmaceutical products, are now appearing in river
394
Although the exchange through the atmosphere of certain trace metals and gases between the land and the sea surface is important, rivers are the main purveyors of materials to the ocean. The total water discharge of the major rivers of the world to the ocean is 36 000 km3 yr 1. At any time, the world’s rivers contain about 0.000 1% of the total water volume of 1459 106 km3 near the surface of the Earth and have a total dissolved and suspended solid concentration of c. 110 and 540 ppm l 1, respectively. The residence time of water in the world’s rivers calculated with respect to total net precipitation on land is only 18 days. Thus the water in the world’s rivers is replaced every 18 days by precipitation. The global annual direct discharge to the ocean of groundwater is about 10% of the surface flow, with a recent estimate of 2400 km3 yr 1. The dissolved constituent content of groundwater is poorly known, but one recent estimate of the dissolved salt groundwater flux is 1300 106 t yr 1. The chemical composition of average river water is shown in Table 1. Note that the major anion in river water is bicarbonate, HCO3 ; the major cation is calcium, Ca2þ , and that even the major constituent concentrations of river water on a global scale are influenced by human activities. The dissolved load of the major constituents of the world’s rivers is derived from the following sources: about 7% from beds of salt and disseminated salt in sedimentary rocks, 10% from gypsum beds and sulfate salts disseminated in rocks, 38% from carbonates, and 45% from the weathering of silicate minerals. Two-thirds of the HCO3 in river waters are derived from atmospheric CO2 via the respiration and decomposition of organic matter and subsequent conversion to HCO3 through the chemical weathering of silicate (B30% of total) or carbonate (B70% of total) minerals. The other third of the river HCO3 comes directly from carbonate minerals undergoing weathering. It is estimated that only about 20% of the world’s drainage basins have pristine water quality. The organic productivity of coastal aquatic environments has been heavily impacted by changes in the
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(c) 2011 Elsevier Inc. All Rights Reserved. 4.6 4.3 1.4 1.4 4.9 4.9 6.7 5.2 3.8 3.8 3.7 3.4 0.3 8%
6.3 6.3
21.2 20.1
31.7 24.2
15.2 15.0
14.7 13.4 1.3 9%
2.2 2.2
17.8 16.6
5.7 5.3
Mg2 þ
7.2 5.2 2.0 28%
7.6 7.0
16.5 3.2
8.4 6.5
3.3 3.3
8.7 6.6
4.4 3.8
Na þ
1.4 1.3 0.1 7%
1.1 1.1
1.8 1.1
1.5 1.5
1.0 1.0
1.7 1.6
1.4 1.4
Kþ
8.3 5.8 2.5 30%
6.8 5.9
20.0 4.7
9.2 7.0
4.1 4.1
10.0 7.6
4.1 3.4
Cl
11.5 5.3 6.2 54%
7.7 6.5
35.5 15.1
18.0 14.9
3.8 3.5
13.3 9.7
4.2 3.2
SO42
53.0 52.0 1.0 2%
65.6 65.1
86.0 80.1
72.3 71.4
24.4 24.4
67.1 66.2
36.9 26.7
HCO 3
10.4 10.4 0.0 0%
16.3 16.3
6.8 6.8
7.2 7.2
10.3 10.3
11.0 11.0
12.0 12.0
SO2
110.1 99.6 10.5 10%
125.3 120.3
212.8 140.3
142.6 133.5
54.6 54.3
134.6 123.5
60.5 57.8
TDSb
12.57 9.89 2.67 21%
TOCb
0.574 0.386 0.187 33%
TDNb
0.053 0.027 0.027 50%
TDPb
37.40 37.40
2.40
2.56
5.53
11.04
12.47
3.41
Water discharge (103 km3 yr 1)
0.46 0.46
0.42
0.38
0.41
0.54
0.28
Runoff ratioc
b
Actual concentrations include pollution. Natural concentrations are corrected for pollution. TDS, total dissolved solids; TOC, total organic carbon; TDN, total dissolved nitrogen; TDP, total dissolved phosphorus. c Runoff ratio ¼ average runoff per unit area/average rainfall. Revised after Meybeck M (1979) Concentrations des eaux fluviales en elements majeurs et apports en solution aux oceans. Revue de Geologie Dynamique et de Geographie Physique 21: 215–246; Meybeck M (1982) Carbon, nitrogen, and phosphorus transport by world rivers. American Journal of Science 282: 401–450; Meybeck M (1983) C, N, P and S in rivers: From sources to global inputs. In: Wollast R, Mackenzie FT, and Chou L (eds.) Interactions of C, N, P and S Biogeochemical Cycles and Global Change, pp. 163–193. Berlin: Springer.
a
Africa Actual Natural Asia Actual Natural S. America Actual Natural N. America Actual Natural Europe Actual Natural Oceania Actual Natural World average Actual Natural (unpolluted) Pollution World % pollutive
Ca2 þ
River water concentrationa (mg l 1)
Chemical composition of average river water
By continent
Table 1
396
LAND–SEA GLOBAL TRANSFERS
dissolved and particulate river fluxes of three of the major bioessential elements found in organic matter, C, N, and P (the other three are S, hydrogen (H), and oxygen (O)). Although these elements are considered minor constituents of river water, their fluxes may have doubled over their pristine values on a global scale because of human activities. Excessive riverborne nutrients and the cultural eutrophication of freshwater and coastal marine ecosystems go hand in hand. In turn, these fluxes have become sensitive indicators of the broader global change issues of population growth and land-use change (including water resources engineering works) in the coastal zone and upland drainage basins, climatic change, and sea level rise. In contrast to the situation for the major elements, delivery of some trace elements from land to the oceans via the atmosphere can rival riverine inputs. The strength of the atmospheric sources strongly depends on geography and meteorology. Hence the North Atlantic, western North Pacific, and Indian Oceans, and their inland seas, are subjected to large atmospheric inputs because of their proximity to both deserts and industrial sources. Crustal dust is the primary terrestrial source of these atmospheric inputs to the ocean. Because of the low solubility of dust in both atmospheric precipitation and seawater and the overwhelming inputs from river sources, dissolved sources of the elements are generally less important. However, because the oceans contain only trace amounts of iron (Fe), aluminum (Al), and manganese (Mn) (concentrations are in the ppb level), even the small amount of dissolution in seawater (B10% of the element in the solid phase) results in eolian dust being the primary source for the dissolved transport of these elements toremote areas of the ocean. Atmospheric transport of the major nutrients N, silicon (Si), and Fe to the ocean has been hypothesized to affect and perhaps limit primary productivity in certain regions of the ocean at certain times. Modern processes of fossil fuel combustion and biomass burning have significantly modified the atmospheric transport from land to the ocean of trace metals like Pb, copper (Cu), and zinc (Zn), C in elemental and organic forms, and nutrient N. As an example of global land–sea transfers involving gases and the effect of human activities on the exchange, consider the behavior of CO2 gas. Prior to human influence on the system, there was a net flux of CO2 out of the ocean owing to organic metabolism (net heterotrophy). This flux was mainly supported by the decay of organic matter produced by phytoplankton in the oceans and part of that transported by rivers to the oceans. An example
overall reaction is: C106 H263 O110 N16 S2 P þ 141O2 ) 106CO2 þ 6HNO3 þ 2H2 SO4 þ H3 PO4 þ 120H2 O
½1
Carbon dioxide was also released to the atmosphere due to the precipitation of carbonate minerals in the oceans. The reaction is: Ca2þ þ 2HCO3 ) CaCO3 þ CO2 þH2 O
½2
The CO2 in both reactions initially entered the dissolved inorganic carbon (DIC) pool of seawater and was subsequently released to the atmosphere at an annual rate of about 0.2 109 t of carbon as CO2 gas. It should be recognized that this is a small number compared with the 200 109 t of carbon that exchanges between the ocean and atmosphere each year because of primary production of organic matter and its subsequent respiration. Despite the maintenance of the net heterotrophic status of the ocean and the continued release of CO2 to the ocean–atmosphere owing to the formation of calcium carbonate in the ocean, the modern ocean and the atmosphere have become net sinks of anthropogenic CO2 from the burning of fossil fuels and the practice of deforestation. Over the past 200 years, as CO2 has accumulated in the atmosphere, the gradient of CO2 concentration across the atmosphere–ocean interface has changed, favoring uptake of anthropogenic CO2 into the ocean. The average oceanic carbon uptake for the decade of the 1990s was c. 2 109 t annually. The waters of the ocean have accumulated about 130 109 t of anthropogenic CO2 over the past 300 years.
The Coastal Zone and Land–Sea Exchange Fluxes The global coastal zone environment is an important depositional and recycling site for terrigenous and oceanic biogenic materials. The past three centuries have been the time of well-documented human activities that have become an important geological factor affecting the continental and oceanic surface environment. In particular, historical increases in the global population in the areas of the major river drainage basins and close to oceanic coastlines have been responsible for increasing changes in land-use practices and discharges of various substances into oceanic coastal waters. As a consequence, the global C cycle and the cycles of N and P that closely interact with the carbon cycle have been greatly affected. Several major perturbations of the past three
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LAND–SEA GLOBAL TRANSFERS
centuries of the industrial age have affected the processes of transport from land, deposition of terrigenous materials, and in situ production of organic matter in coastal zone environments. In addition, potential future changes in oceanic circulation may have significant effects on the biogeochemistry and CO2 exchange in the coastal zone. The coastal zone is that environment of continental shelves, including bays, lagoons, estuaries, and near-shore banks that occupy 7% of the surface area of the ocean (36 1012 m2) or c. 9% of the volume of the surface mixed layer of the ocean (3 1015 m3). The continental shelves average 75 km in width, with a bottom slope of 1.7 m km 1. They are generally viewed as divisible into the interior or proximal shelf, and the exterior or distal shelf. The mean depth of the global continental shelf is usually taken as the depth of the break between the continental shelf and slope at c. 200 m, although this depth varies considerably throughout the world’s oceans. In the Atlantic, the median depth of the shelf-slope break is at 120 m, with a range from 80 to 180 m. The depths of the continental shelf are near 200 m in the European section of the Atlantic, but they are close to 100 m on the African and North American coasts. Coastal zone environments that have high sedimentation rates, as great as 30–60 cm ky–1 in active depositional areas, act as traps and filters of natural and human-generated materials transported from continents to the oceans via river and groundwater flows and through the atmosphere. At present a large fraction (B80%) of the land-derived organic and inorganic materials that are transported to the oceans is trapped on the proximal continental shelves. The coastal zone also accounts for 30–50% of total carbonate and 80% of organic carbon accumulation in the ocean. Coastal zone environments are also regions of higher biological production relative to that of average oceanic surface waters, making them an important factor in the global carbon cycle. The higher primary production is variably attributable to the nutrient inflows from land as well as from coastal upwelling of deeper ocean waters. Fluvial and atmospheric transport links the coastal zone to the land; gas exchange and deposition are its links with the atmosphere; net advective transport of water, dissolved solids and particles, and coastal upwelling connect it with the open ocean. In addition, coastal marine sediments are repositories for much of the material delivered to the coastal zone. In the last several centuries, human activities on land have become a geologically important agent affecting the land–sea exchange of materials. In particular, river and groundwater flows and atmospheric
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transport of materials to the coastal zone have been substantially altered. Bioessential Elements
Continuous increase in the global population and its industrial and agricultural activities have created four major perturbations on the coupled system of the biogeochemical cycles of the bioessential elements C, N, P, and S. These changes have led to major alterations in the exchanges of these elements between the land and sea. The perturbations are: (1) emissions of C, N, and S to the atmosphere from fossil-fuel burning and subsequent partitioning of anthropogenic C and deposition of N and S; (2) changes in land-use practices that affect the recycling of C, N, P, and S on land, their uptake by plants and release from decaying organic matter, and the rates of land surface denudation; (3) additions of N and P in chemical fertilizers to cultivated land area; and (4) releases of organic wastes containing highly reactive C, N, and P that ultimately enter the coastal zone. A fifth major perturbation is a climatic one: (5) the rise in mean global temperature of the lower atmosphere of about 1 1C in the past 300 years, with a projected increase of about 1.4–5.8 1C relative to 1990 by the year 2100. Figure 1 shows how the fluxes associated with these activities have changed during the past three centuries with projections to the year 2040. Partially as a result of these activities on land, the fluxes of materials to the coastal zone have changed historically. Figure 2 shows the historical and projected future changes in the river fluxes of dissolved inorganic and organic carbon (DIC, DOC), nitrogen (DIN, DON), and phosphorus (DIP, DOP), and fluxes associated with the atmospheric deposition and denitrification of N, and accumulation of C in organic matter in coastal marine sediments. It can be seen in Figure 2 that the riverine fluxes of C, N, and P all increase in the dissolved inorganic and organic phases from about 1850 projected to 2040. For example, for carbon, the total flux (organic þ inorganic) increases by about 35% during this period. These increased fluxes are mainly due to changes in land-use practices, including deforestation, conversion of forest to grassland, pastureland, and urban centers, and regrowth of forests, and application of fertilizers to croplands and the subsequent leaching of N and P into aquatic systems. Inputs of nutrient N and P to the coastal zone which support new primary production are from the land by riverine and groundwater flows, from the open ocean by coastal upwelling and onwelling, and to a lesser extent by atmospheric deposition of nitrogen. New primary production depends on the
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LAND–SEA GLOBAL TRANSFERS
(d) 150
10
Fossil fuel
Carbon
100
5
50 Nitrogen Land use
0
0
Phosphorus
Sewage input (Mt yr −1)
C emissions (Gt yr −1)
(a)
(e)
100 Sulfur
50 0
Nitrogen
N, S emissions (Mt yr −1)
150
1.5 Δ temperature (°C)
(b)
1.0 0.5 0
Fertilizer input (Mt yr −1)
(c) 1700
150
1800
1900 Year
2000
Nitrogen
100 50 0 1700
Phosphorus
1800
1900 Year
2000
Figure 1 Major perturbations on the Earth system over the past 300 years and projections for the future: (a) emissions of CO2 and (b) gaseous N and S from fossil-fuel burning and land-use activities; (c) application of inorganic N and P in chemical fertilizers to cultivated land; (d) loading of highly reactive C, N, and P into rivers and the coastal ocean from municipal sewage and wastewater disposal; and (e) rise in mean global temperature of the lower atmosphere relative to 1700. Revised after Ver LM, Mackenzie FT, and Lerman A (1999) Biogeochemical responses of the carbon cycle to natural and human perturbations: Past, present, and future. American Journal of Science 299: 762–801.
availability of nutrients from these external inputs, without consideration of internal recycling of nutrients. Thus any changes in the supply of nutrients to the coastal zone owing to changes in the magnitude of these source fluxes are likely to affect the cycling pathways and balances of the nutrient elements. In particular, input of nutrients from the open ocean by coastal upwelling is quantitatively greater than the combined inputs from land and the atmosphere. This makes it likely that there could be significant effects on coastal primary production because of changes in ocean circulation. For example, because of global warming, the oceans could become more strongly stratified owing to freshening of polar oceanic waters and warming of the ocean in the tropical zone. This could lead to a reduction in the intensity of the oceanic thermohaline circulation (oceanic circulation owing to differences in density of water masses, also
popularly known as the ‘conveyor belt’) and hence the rate at which nutrient-rich waters upwell into coastal environments. Another potential consequence of the reduction in the rate of nutrient inputs to the coastal zone by upwelling is the change in the CO2 balance of coastal waters: reduction in the input of DIC to the coastal zone from the deeper ocean means less dissolved CO2, HCO3 , and CO32 coming from that source. With increasing accumulation of anthropogenic CO2 in the atmosphere, the increased dissolution of atmospheric CO2 in coastal water is favored. The combined result of a decrease in the upwelling flux of DIC and an enhancement in the transfer of atmospheric CO2 across the air–sea interface of coastal waters is a lower saturation state for coastal waters with respect to the carbonate minerals calcite, aragonite, and a variety of magnesian calcites. The
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LAND–SEA GLOBAL TRANSFERS
(a) 50
Carbon
Riverine total organic flux 40
Riverine dissolved inorganic flux
30
Organic carbon Accumulation in sediments 20
10 (b) 6
Nitrogen Riverine total organic flux
5 4
Denitrification
3 2
399
lower saturation state in turn leads to the likelihood of lower rates of inorganic and biological precipitation of carbonate and hence deposition and accumulation of sedimentary carbonate. In addition, the present-day burial rate of organic carbon in the ocean may be about double that of the late Holocene flux, supported by increased fluxes of organic carbon to the ocean via rivers and groundwater flows and increased in situ new primary production supported by increased inputs of inorganic N and P from land and of N deposited from the atmosphere. The organic carbon flux into sediments may constitute a sink of anthropogenic CO2 and a minor negative feedback on accumulation of CO2 in the atmosphere. The increased flux of land-derived organic carbon delivered to the ocean by rivers may accumulate there or be respired, with subsequent emission of CO2 back to the atmosphere. This release flux of CO2 may be great enough to offset the increased burial flux of organic carbon to the seafloor due to enhanced fertilization of the ocean by nutrients derived from human activities. The magnitude of the CO2 exchange is a poorly constrained flux today. One area for which there is a substantial lack of knowledge is the Asian Pacific region. This is an area of several large seas, a region of important river inputs to the ocean of N, P, organic carbon, and sediments from land, and a region of important CO2 exchange between the ocean and the atmosphere.
Atmospheric deposition flux
Anticipated Response to Global Warming 1 Riverine dissolved inorganic flux
0 (c)
Phosphorus 0.6
0.4 Riverine total organic flux 0.2 Riverine dissolved inorganic flux 0 1850
1950
1900
2000
Year
Figure 2 Past, present, and predicted fluxes of carbon, nitrogen, and phosphorus into or out of the global coastal margin, in 1012 mol yr 1.
From 1850 to modern times, the direction of the net flux of CO2 between coastal zone waters and the atmosphere due to organic metabolism and calcium carbonate accumulation in coastal marine sediments was from the coastal surface ocean to the atmosphere (negative flux, Figure 3).This flux in 1850 was on the order of 0.2 109 t yr 1. In a condition not disturbed by changes in the stratification and thermohaline circulation of the ocean brought about by a global warming of the Earth, the direction of this flux is projected to remain negative (flux out of the coastal ocean to the atmosphere) until early in the twenty-first century. The increasing partial pressure of CO2 in the atmosphere because of emissions from anthropogenic sources leads to a reversal in the gradient of CO2 across the air–sea interface of coastal zone waters and, hence, invasion of CO2 into the coastal waters. From that time on the coastal ocean will begin to operate as a net sink (positive flux) of atmospheric CO2 (Figure 3). The role of the open ocean as a sink for anthropogenic CO2 is slightly reduced while that of the coastal oceans
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LAND–SEA GLOBAL TRANSFERS
5.0 No change 34% reduction 50% reduction 100% reduction
160
6.0
4.0
5.5
Calcite
120
2.0
40
1.0
(+) flux from atmosphere 0
0
to coastal waters
Carbonate saturation state, Ω
80
(Mol C m−2yr−1)
(1012 mol C yr−1)
3.0
5.0
4.5 Aragonite 4.0
3.5 _1.0
1850
1900
1950 Year
2000
15 mol.% magnesian calcite
3.0 Figure 3 The net flux of CO2 between coastal zone waters and the atmosphere due to organic metabolism and calcium carbonate accumulation in coastal marine sediments, under three scenarios of changing thermohaline circulation rate compared to a business-as-usual scenario, in units of 1012 mol C yr 1.
increases. The net result is the maintenance of the role of the global oceans as a net sink for anthropogenic CO2. The saturation index (O) for calcite or aragonite (both CaCO3) is the ratio of the ion activity product IAP in coastal waters to the respective equilibrium constant K at the in situ temperature. For aqueous species, IAP ¼ aCa2þ aCO32 (where a is the activity; note that for 15 mol.% magnesian calcite, the IAP also includes the activity of the magnesium cation, aMg2þ ). Most coastal waters and open-ocean surface waters currently are supersaturated with respect to aragonite, calcite, and magnesian calcite containing 15 mol.% Mg, that is, Ocalcite, Oaragonite, and O15%magnesian calcite are 41. Because of global warming and the increasing land-to-atmosphere-toseawater transport of CO2 (due to the continuing combustion of fossil fuels and burning of biomass), the concentration of the aqueous CO2 species in seawater increases and the pH of the water decreases slightly. This results in a decrease in the concentration of the carbonate ion, CO322 , resulting in a decrease in the degree of supersaturation of coastal zone waters. Figure 4 shows how the degree of saturation might change into the next century because of rising atmospheric CO2 concentrations. The overall reduction in the saturation state of coastal
1980
2000
2020
2040
Year Figure 4 Changes in saturation state with respect to carbonate minerals of surface waters of the coastal ocean projected from 1999 to 2040. Calculations are for a temperature of 25 1C.
zone waters with respect to aragonite from 1997 projected to 2040 is about 16%, from 3.89 to 3.26. Modern carbonate sediments deposited in shoalwater (‘shallow-water’) marine environments (including shelves, banks, lagoons, and coral reef tracts) are predominantly biogenic in origin derived from the skeletons and tests of benthic and pelagic organisms, such as corals, foraminifera, echinoids, mollusks, algae, and sponges. One exception to this statement is some aragonitic muds that may, at least in part, result from the abiotic precipitation of aragonite from seawater in the form of whitings. Another exception is the sand-sized, carbonate oo¨ids composed of either aragonite with a laminated internal structure or magnesian calcite with a radial internal structure. In addition, early diagenetic carbonate cements found in shoal-water marine sediments and in reefs are principally aragonite or magnesian calcite. Thus carbonate production and accumulation in shoal-water environments are dominated by a range of metastable carbonate minerals associated with skeletogenesis and abiotic processes, including calcite, aragonite, and a variety of magnesian calcite compositions.
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LAND–SEA GLOBAL TRANSFERS
With little doubt, as has been documented in a number of observational and experimental studies, a reduction in the saturation state of ocean waters will lead to a reduction in the rate of precipitation of both inorganic and skeletal calcium carbonate. Conversely, increases in the degree of supersaturation and temperature will increase the precipitation rates of calcite and aragonite from seawater. During global warming, rising sea surface temperatures and declining carbonate saturation states due to the absorption of anthropogenic CO2 by surface ocean waters result in opposing effects. However, experimental evidence suggests that within the range of temperature change predicted for the next century due to global warming, the effect of changes in saturation state will be the predominant factor affecting precipitation rate. Thus decreases in precipitation rates should lead to a decrease in the production and accumulation of shallow-water carbonate sediments and perhaps changes in the types and distribution of calcifying biotic species found in shallow-water environments.
Anticipated Response to Heightened Human Perturbation: The Asian Scenario In the preceding sections it was shown that the fluxes of C, N, and P from land to ocean have increased because of human activities (refer to Table 1 for data comparing the actual and natural concentrations of C, N, P, and other elements in average river water). During the industrial era, these fluxes mainly had their origin in the present industrialized and developed countries. This is changing as the industrializing and developing countries move into the twenty-first century. A case in point is the countries of Asia. Asia is a continent of potentially increasing contributions to the loading of the environment owing to a combination of such factors as its increasing population, increasing industrialization dependent on fossil fuels, concentration of its population along the major river drainage basins and in coastal urban centers, and expansion of land-use practices. It is anticipated that Asia will experience similar, possibly even greater, loss of storage of C and nutrient N and P on land and increased storage in coastal marine sediments per unit area than was shown by the developed countries during their period of industrialization. The relatively rapid growth of Asia’s population along the oceanic coastal zone indicates that higher inputs of both dissolved and particulate organic nutrients may be expected to enter coastal waters.
401
A similar trend of increasing population concentration in agricultural areas inland, within the drainage basins of the main rivers flowing into the ocean, is also expected to result in increased dissolved and particulate organic nutrient loads that may eventually reach the ocean. Inputs from inland regions to the ocean would be relatively more important if no entrapment or depositional storage occurred en route, such as in the dammed sections of rivers or in alluvial plains. In the case of many of China’s rivers, the decline in sediment discharge from large rivers such as the Yangtze and the Yellow Rivers is expected to continue due to the increased construction of dams. The average decadal sediment discharge from the Yellow River, for example, has decreased by 50% from the 1950s to the 1980s. If the evidence proposed for the continental United States applies to Asia, the damming of major rivers would not effectively reduce the suspended material flow to the ocean because of the changes in the erosional patterns on land that accompany river damming and more intensive land-use practices. These flows on land and into coastal ocean waters are contributing factors to the relative importance of autotrophic and heterotrophic processes, competition between the two, and the consequences for carbon exchange between the atmosphere and land, and the atmosphere and ocean water. The change from the practices of land fertilization by manure to the more recent usage of chemical fertilizers in Asia suggests a shift away from solid organic nutrients and therefore a reduced flow of materials that might promote heterotrophy in coastal environments. Sulfur is an excellent example of how parts of Asia can play an important role in changing land–sea transfers of materials. Prior to extensive human interference in the global cycle of sulfur, biogenically produced sulfur was emitted from the sea surface mainly in the form of the reduced gas dimethyl sulfide (DMS). DMS was the major global natural source of sulfur for the atmosphere, excluding sulfur in sea salt and soil dust. Some of this gas traveled far from its source of origin. During transport the reduced gas was oxidized to micrometer-size sulfate aerosol particles and rained out of the atmosphere onto the sea and continental surface. The global sulfur cycle has been dramatically perturbed by the industrial and biomass burning activities of human society. The flux of gaseous sulfur dioxide to the atmosphere from the combustion of fossil fuels in some regions of the world and its conversion to sulfate aerosol greatly exceeds natural fluxes of sulfur gases from the land surface. It is estimated that this flux for the year 1990 was equivalent to 73 106 t yr1, nearly 4 times the natural DMS flux from the
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ocean. This has led to a net transport of sulfur from the land to the ocean via the atmosphere, completely reversing the flow direction in preindustrial times. In addition, the sulfate aerosol content of the atmosphere derived from human activities has increased. Sulfate aerosols affect global climate directly as
particles that scatter incoming solar radiation and indirectly as cloud condensation nuclei (CCNs), which lead to an increased number of cloud droplets and an increase in the solar reflectance of clouds. Both effects cause the cooling of the planetary surface. As can be seen in Figure 5 the eastern Asian
(a)
2500 1000 500 250
(b)
100 50 10
Figure 5 Comparison of the magnitude of atmospheric sulfur deposition for the years 1990 (a) and 2050 (b). Note the large increases in both spatial extent and intensity of sulfur deposition in both hemispheres and the increase in importance of Asia, Africa, and South America as sites of sulfur deposition between 1990 and 2050. The values on the diagrams are in units of kg S m 2 yr 1. Revised after Mackenzie FT (1998) Our Changing Planet : An Introduction to Earth System Science and Global Environmental Change. Upper Saddle River, NJ: Prentice Hall; Rodhe H, Langner J, Gallardo L, and Kjellstro¨m E (1995) Global transport of acidifying pollutants. Water, Air and Soil Pollution 85: 37–50.
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LAND–SEA GLOBAL TRANSFERS
region is an important regional source of sulfate aerosol because of the combustion of fossil fuels, particularly coal. This source is predicted to grow in strength during the early- to mid-twenty-first century (Figure 5).
Conclusion Land–sea exchange processes and fluxes of the bioessential elements are critical to life. In several cases documented above, these exchanges have been substantially modified by human activities. These modifications have led to a number of environmental issues including global warming, acid deposition, excess atmospheric nitrogen deposition, and production of photochemical smog. All these issues have consequences for the biosphere – some well known, others not so well known. It is likely that the developing world, with increasing population pressure and industrial development and with no major changes in agricultural technology and energy consumption rates, will become a more important source of airborne gases and aerosols and materials for river and groundwater systems in the future. This will lead to further modification of land–sea global transfers. The region of southern and eastern Asia is particularly well poised to influence significantly these global transfers.
See also Aeolian Inputs. Air–Sea Transfer: Dimethyl Sulfide, COS, CS2, NH4, Non-Methane Hydrocarbons, Organo-Halogens. Carbon Cycle. Carbon Dioxide (CO2) Cycle. Coastal Topography, Human Impact on. Nitrogen Cycle. Ocean Carbon System, Modeling of. Ocean Circulation: Meridional Overturning Circulation. Past Climate from Corals. Phosphorus Cycle.
403
Further Reading Berner EA and Berner RA (1996) Global Environment: Water, Air and Geochemical Cycles. Upper Saddle River, NJ: Prentice Hall. Galloway JN and Melillo JM (eds.) (1998) Asian Change in the Context of Global Change. Cambridge, MA: Cambridge University Press. Mackenzie FT (1998) Our Changing Planet: An Introduction to Earth System Science and Global Environmental Change. Upper Saddle River, NJ: Prentice Hall. Mackenzie FT and Lerman A (2006) Carbon in the Geobiosphere – Earth’s Outer Shell. Dordrecht: Springer. Meybeck M (1979) Concentrations des eaux fluviales en elements majeurs et apports en solution aux oceans. Revue de Geologie Dynamique et de Geographie Physique 21: 215--246. Meybeck M (1982) Carbon, nitrogen, and phosphorus transport by world rivers. American Journal of Science 282: 401--450. Meybeck M (1983) C, N, P and S in rivers: From sources to global inputs. In: Wollast R, Mackenzie FT, and Chou L (eds.) Interactions of C, N, P and S Biogeochemical Cycles and Global Change, pp. 163--193. Berlin: Springer. Rodhe H, Langner J, Gallardo L, and Kjellstro¨m E (1995) Global transport of acidifying pollutants. Water, Air and Soil Pollution 85: 37--50. Schlesinger WH (1997) Biogeochemistry: An Analysis of Global Change. San Diego, CA: Academic Press. Smith SV and Mackenzie FT (1987) The ocean as a net heterotrophic system: Implications from the carbon biogeochemical cycle. Global Biogeochemical Cycles 1: 187--198. Ver LM, Mackenzie FT, and Lerman A (1999) Biogeochemical responses of the carbon cycle to natural and human perturbations: Past, present, and future. American Journal of Science 299: 762--801. Vitousek PM, Aber JD, and Howarth RW (1997) Human alteration of the global nitrogen cycle: Sources and consequences. Ecological Applications 7(3): 737--750. Wollast R and Mackenzie FT (1989) Global biogeochemical cycles and climate. In: Berger A, Schneider S, and Duplessy JC (eds.) Climate and Geo-Sciences, pp. 453--473. Dordrecht: Kluwer Academic Publishers.
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LANGMUIR CIRCULATION AND INSTABILITY S. Leibovich, Cornell University, Ithaca, NY, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The surface of a wind-driven sea often is marked by streaks roughly aligned with the wind direction. These streaks, or windrows, are visible manifestations of coherent subsurface motions extending throughout the bulk of the ocean surface mixed layer, extending from the surface down to the seasonal thermocline. These may be regarded as the large scales of the turbulence in the mixed layer. Windrows and their subsurface origins were first systematically studied and described by Irving Langmuir in 1938, and the phenomenon since has become known as Langmuir circulation. The existence of a simple deterministic description making these large scales theoretically accessible distinguishes this problem from coherent structures in other turbulent flows. The theory traces these patterns to a convective instability mechanically driven by the wind waves and currents. Recent advances in instrumentation and computational data analysis have led to field observations of Langmuir circulation of unprecedented detail. Although the body of observational data obtained since Langmuir’s own work is mainly qualitative, ocean experiments now can yield quantitative measurements of velocity fields in the near-surface region. New measurement methods are capable of producing data comprehensive enough to characterize the phenomenon, and its effect on the stirring and maintenance of the mixed layer, although the labor and difficulties involved and the shear complexity of the processes occurring in the surface layer leave much work to be done before this can be said to be accomplished. Nevertheless, the combination of new experimental techniques and a simple and testable theoretical mechanism has stimulated rapid progress in the exploration of the stirring of the ocean surface mixed layer.
mechanical processes through the action of the wind, as Langmuir originally indicated. At the surface, rolls act to sweep surface water from regions of surface divergence overlying upwelling water into convergence zones overlying downwelling water. Floating material is collected into lines of surface convergence visible as windrows. In confined bodies of water, such as lakes and ponds, windrows are very nearly parallel to the wind, and can have a nearly uniform spacing as shown in Figure 2. In the open ocean, evidence indicates windrows tend to be oriented at small angles to the wind (typically to the right in the Northern Hemisphere), spacing is more variable, and individual windrows can be traced only for a modest multiple of the mean spacing. A windrow may either terminate, perhaps due to local absence of surface tracers, coalesce with an adjacent windrow, or split into two daughter windrows. Thus in the ocean, the general surface appearance is of a network of lines, occasionally interecting, yet roughly aligned with the wind. Windrows are visible in nature only when both Langmuir circulation and surface tracers are present. In the ocean, bubbles from breaking waves are the most readily available tracers, and Langmuir
z
Wind <15°
Co ak re St
ak re e St on z
ce
en
g er
v
Di
w _ 1_ 2 cm s 1
x
y
e nc ge ne r e zo nv
e
2_ 3 _ cm s 1
c en
rg
u ~10 relativ divergence z
ve
n Co
w
2_ 6 cm s
_1
Ro ws
pa ci
ng (5 _ 30
0m )
Description of Langmuir Circulation Langmuir circulation takes the ideal form of vortices with axes aligned the wind, as in the schematic drawing in Figure 1. The appearance resembles convective rolls driven by thermal convection, but all evidence indicates that the motions are due to
404
Figure 1 Sketch of Langmuir circulation. From Pollard RT (1977) Observations and theories of Langmuir circulations and their role in near surface mixing. In: Angel M (ed.) A Voyage of Discovery: G. Deacon 70th Anniversary Volume, Re, pp. 235– 251. Oxford, UK: Pergamon.
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LANGMUIR CIRCULATION AND INSTABILITY
Figure 2 Photograph of windrows in Rodeo Lagoon, California. From Szeri AJ (1996) Langmuir circulations on Rodeo Lagoon. Monthly Weather Review 124(2): 341–342.
Figure 3 Infrared aerial photograph of thermal streaks in the Pacific. The arrow shows the wind direction. The wind speed is 4 m s 1. The image is about 750 m wide. From McLeish W (1968) On the mechanisms of wind-slick generation. Deep-Sea Research 15: 461–469.
circulation and bubbles both appear to exist when speeds exceed some threshold. Threshold wind levels are not absolute, since swell, wind duration, fetch, and currents existing before the onset of wind forcing play a role, but windrows are commonly reported in winds of 3 m s 1 or more. Tracers other than bubbles may produce windrows revealing underlying Langmuir circulation – all forms of flotsam serve. Langmuir first noticed windrows from the deck of a ship in the Sargasso Sea, which contained windrows of Sargassum. Organic films on the water surface are compressed in windrows, causing capillary waves to be preferentially damped; in light winds, windrows thereby are made visible as bands of smoother water. Observations in the infrared (see Figure 3) reveal windrows due to variation of surface temperature created by Langmuir circulation.
405
A hierarchy of horizontal scales is observed, with windrow spacing ranging in the ocean from a few meters up to approximately 3 times the depth of the surface mixed layer. The largest scales are the most energetic and the most persistent, and extend to depths comparable to that of the mixed layer, so a cell extending from surface divergence to surface convergence, and from the surface to maximum depth of penetration, is approximately square. The maximum penetration depth in the open ocean is comparable to the depth of the seasonal thermocline. Langmuir reported the penetration depth in Lake George to be comparable to the epilimnion when the lake was stratified, and believed the epilimnion to be created by the mixing caused by the wind-driven convective motion. Whether the seasonal thermocline location is fixed by the scale of Langmuir circulation, or whether the penetration depth of the circulation is limited by the strong buoyancy at the thermocline acting like a bottom is not yet clear, although the latter seems more likely. In shallow water, the seabed or lake floor of course fixes the maximum penetration depth. While horizontal scales of up to 3 times the mixed layer depth are reported, it is not clear that even larger scales exist that have not been detected in experiments. The smaller scales are advected and presumably eventually swept up by the largest scales. If a fixed number of permanent surface markers were used, as in some experiments where computer cards were released to serve as markers, the larger scales ultimately would be more prevalent. Most observations do depend on Lagrangian tracers, in particular bubbles with definite lifetimes that are continually but episodically created. The regeneration of bubbles on the surface between the large-scale windrows permits the smaller scales to be seen. The largest observed windrow scales consolidate in about 20 min after a shift in wind direction, or in cases of sudden wind onset. This appears to establish the formation time, at least that required to sweep surface material into windrows. The largest windrow scales in the Pacific have been observed to persist for hours. Downwelling speeds below windrows are substantially higher than upwelling speeds occurring below surface divergences. Furthermore, the downwind surface speed is larger in windrows than between them. A simple explanation for the speed excess is given in the section on instability. Although the observed speed increase is commonly reported, it has not often been quantified. It appears, however, that the speed increases in surface jets seem to be comparable to the maximum downwelling speeds.
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Theory Theory promulgated in the 1970s has influenced experiments addressing Langmuir circulation. The theory most commonly utilized, and now commonly referred to as the Craik–Leibovich theory after its originators, begins with Langmuir’s conclusion that the cellular motion bearing his name derives from the wind. Wind blowing over a water surface has two simultaneous consequences: currents are generated as horizontal momentum is transmitted from wind to water; and waves are generated on the water surface due to an instability of the air–sea interface under wind shear. A detailed treatment of the shear flow in the presence of wind waves is not feasible. The eddy turnover timescale (time required for a fluid particle to traverse the convective cell) in Langmuir circulation is on the order of tens of minutes. Surface waves have a substantially shorter timescale. Winddriven water waves can be thought of as comprised of the superposition of wavelets with a continuous range of wavelengths and frequencies, and the amplitudes of waves in a given band of wavelength or frequency can be characterized by an energy spectrum. For the Pierson–Moskowitz empirical wind wave spectrum, the waves at the peak of the spectrum of a wind-driven sea under wind speed U have period approximately 7g/U. This is a typical value for the energetic part of the wind wave spectrum. For wind speeds leading to observable Langmuir circulation, a characteristic peak wave period is of the order of 10 s. Averaging over the waves is therefore useful. Surface gravity waves are approximately irrotational, and orbital speeds near the surface are an order of magnitude larger than the mean speeds in the current system. Although a wind-driven wave field is complex, and generally must be treated probabilistically, the theory depends only on averaged effects due to the net mass drift caused by the waves, and this often can be computed. Wave Effects: The Stokes Drift
Nonbreaking surface gravity waves have small slopes, and produce a small (second order in wave slope) mean mass motion in the direction of wave propagation. This net water motion is known as the Stokes drift. If uw is the Eulerian velocity due to the water waves, the Stokes drift velocity us is approximately given by us ¼
Z
t
uw dt ruw
½1
where the angle brackets represent an average over time. An acceptable averaging time to calculate
Stokes drift is a modest multiple of the wave period associated with the peak frequency of the energy spectrum. Provided the currents are small compared to the wave orbital speeds, as is usually assumed in the ocean, the net mass drift due to waves can be approximated by this formula. If the surface wave characteristics can be determined with specified wind fields – for example, if the wave spectrum is measured, or computed by a wave forecasting model, or if it is assumed to follow a standard empirical wave spectrum – then the Stokes drift can be calculated using [1]. The form of the vertical variation will depend on the spectrum, and the resulting functional form will depend on two parameters (the surface value of the drift, Us, and a length scale, c), representing the characteristic decay depth of the drift. For example, for a small-amplitude monochromatic surface gravity wave with wavelength l and amplitude, a, the Stokes drift velocity is given by jus j ¼ Us expðx3 =cÞ;
2pa2 Us ¼ l
rffiffiffiffiffiffiffiffi 2pg ; l
c¼
l 4p ½2
where x3 is the coordinate measured vertically upward from the mean free surface. In the absence of swell, the direction of the drift is parallel to the wind. Langmuir Force
Assuming the surface waves can be approximated as irrotational and have orbital speeds large compared to current speeds, averaging yields an apparent extra force on the averaged motion equal to us xa, where us is the Stokes drift velocity associated with the wave motion, and xa is the averaged absolute vorticity of the water body. The apparent extra force captures mean effects of the wave–current interaction. It has been called the vortex force in the literature, although it might be more appropriate to call it the Langmuir force, which is the term adopted here. The mean Lagrangian velocity, or mass drift, of a fluid particle at position x at time t exceeds the mean Eulerian velocity of the water by us. If the system is referred to a frame rotating with constant angular velocity vector X, then the Langmuir force FL is FL ðx; tÞ ¼ us ðx; tÞ ð2X þ xðx; tÞÞ
½3
where x ¼ curl u. Note that this quantity has dimensions of a force per unit mass. Here u is the Eulerian mean velocity as seen in the rotating frame, and so x is the relative mean vorticity. The full instantaneous velocity field includes fluctuations due
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LANGMUIR CIRCULATION AND INSTABILITY
turbulence as well as waves, and so the waveaveraged mean velocity filters out not only the waves, but also turbulence with timescales comparable to or less than the averaging time. The Boussinesq equations for momentum and temperature, when averaged over a timescale of several wave periods, are @u 1 ¯ 3 þ u ru þ 2X u ¼ rp þ bgðT TÞe @t r þ us ð2X þ xÞ þ F
½4
@T þ ðu þ us Þ rT ¼ H @t
½5
ru¼0
½6
In these equations, p is a modified pressure including mean wave effects, centripetal force ‘potential’, and hydrostatic pressure variations; T is the water temperature; T¯ is a reference temperature field that may depend on time and depth; e3 is a unit vector pointing vertically upward; while F and H represent the divergence of momentum and heat fluxes, respectively, due to the unresolved turbulent scales. The rectified effects of the waves are represented by the Stokes drift in the Langmuir force and in the advection of the scala T. If other scalar quantities, like salt concentration, were to be included, they would be governed by equations of the same form as that for T. With suitable subgrid models for F and H, these equations can be used to compute the motions, including those with turbulent fluctuations having timescales large compared to a wave period. The principal theoretical development is the incorporation of residual wave effects the Langmuir force and scalar advection augmented by the Stokes drift velocity. Without these additions, the governing equations are the same as conventional models not accounting for surface waves. The simplest closure models are the assumptions of constant eddy viscosity, nT , and eddy diffusity of heat, kT , yielding F ¼ nT r2 u and H ¼ kT r2 T. All analytical work done addressing the stability of the mean motion assumes this form, as does much of the computational work on Langmuir circulation. Large eddy simulations (LESs) of turbulent Langmuir circulation using standard subgrid closure models for F and H have been carried out by Skyllingstad and Denbo and by McWilliams et al., among others. Calculations of this sort, while involving a great deal of computational effort, have the important advantage of dispensing with empirical eddy coefficients nT , and kT. The relative effects of the Langmuir force and ordinary shear turbulence can be
407
educed from LES using the ‘turbulent Langmuir pffiffiffiffiffiffiffiffiffiffiffiffiffi number’ defined by McWilliams et al. to be u =Us . Here u is the friction velocity defined in the next section. The discussion so far has been predicated on the existence of wind stress to produce currents and surface waves. Currents due to other causes may exist for a time in the absence of wind. The same is true of surface waves, which may have been generated in distant locations and propagate to the region of interest. The Langmuir force continues to exist, and under suitable conditions – especially when the angle between the Stokes drift and current is small – Langmuir circulation may result. Thus, the details of the current and wave fields must be known, and scalings based solely on wind or Stokes drift maybe misleading. Scaling
Despite the disclaimer of the previous section, reported Langmuir circulation appears to be confined to situations driven by local wind stress. The motion will depend on the magnitude and (if Coriolis effects apply) direction of the applied wind stress, s sea state as prescribed by us; the prevailing density stratification; surface thermal conditions; water depth, or specification of conditions of current, temperature, and salinity below the mixed layer; the Coriolis acceleration, which depends on latitude; and initial conditions. The wind stress clearly is a primary factor; itpprovides ffiffiffiffiffiffiffiffiffiffi a unit for speed, the friction velocity u ¼ jsj=r, r being the water density. Several length scales appear, and the choice of unit of length, d, will depend on circumstances. Since observation in the ocean and in stratified lakes indicates the importance of the depth of the seasonal thermocline, it provides a natural choice when it can be defined. Other choices replace this when a strong thermocline does not exist. Clearly, the range of questions that may be encountered is extensive, and the parameter space encompassing them is too large for a general discussion. The most elementary situation producing Langmuir circulation, however, is that of a nonrotating layer of water of uniform density with finite depth (either to a solid bottom or to strong thermocline that strongly inhibits vertical motion) under the action of a wind of unlimited fetch and duration and with constant speed and direction producing a sea state with known Stokes drift. With u as unit for speed, d unit for length, the unit for time is d=u , and the problem then depends on the dimensionless parameters u =Us and c=d, and a parameter characterizing Fd=u2 . The latter depends on how F is modeled. If a constant
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eddy viscosity is used, then the appropriate parameter is Re1 , where Re ¼ u d=nT is the Reynolds number based on eddy viscosity. The designation Re can serve as a placeholder for the parameter appropriate for the choice of model. Thus, the simplest problem depends on three dimensionless parameters, although if the motion is restricted to be invariant in the wind direction, the parameter space can be reduced by one. For each additional physical effect, one more parameter is added to the list. For example, if rotation effects are added, the parameter jXjd=u, an inverse Rossby number, is needed. For the most elementary case, the velocity vector in the water column is given by u ¼ u fðx=d; t d=u ; Us =u ; c=d; Re Þ, where f is a dimensionless vector-valued function of its arguments. For fixed position and time, the motion depends on the three fixed dimensionless parameters. If the dependence on Re and c=d were weak, hypotheses about how f depends on Us =u suggest ways to scale experimental data. For example, if the dependence is linear, then u is directly proportional to Us. If the dependence is on the square root, then u is directly proportional to pffiffiffiffiffiffiffiffiffiffi u Us . Both of these scalings have been tried for field data, without conclusive results. In fact, constant eddy viscosity simulations using [4] and [6] fail to indicate a simple functional dependence of f on its parameters.
Instability
If all motions are due to wind and perhaps buoyancy forcing and the wind stress, sea state, and thermal boundary conditions are independent of horizontal position, the theoretical model has exact solutions independent of horizontal position. The natural mechanical condition to impose in such cases is that of a constant surface stress in a given direction, and, if waves are present, a Stokes drift velocity parallel to the applied stress. These solutions describe flows with no vertical motion and no patterns on the surface. Such flows therefore may be described as featureless. These exact solutions can be unstable, however, and patterns can develop as a consequence. In the absence of wave effects us and FL both vanish, and roll instabilities leading to parallel lines of surface convergence can arise in two physically distinct ways. If the Coriolis acceleration is retained and thermal effects are ignored, the featureless flow is the wellknown stress-driven Ekman layer, and it is unstable for sufficiently large wind stress. (The same problem without Coriolis effects yields plane Couette flow, which is stable.) If thermal effects are accounted for
and the water is cooled from above, the motion is unstable for sufficiently large cooling rates. In the absence of both Coriolis and thermal effects, the motion without surface waves is stable for any value of the surface stress. If waves are present, the action of the Langmuir force causes the flow to be unstable when the applied stress exceeds a threshold value. The preferred instability mode consists of rolls parallel to the wind direction. If Coriolis accelerations and Langmuir force are simultaneously considered for typical oceanic conditions, the flow instability s close to that found for nonrotating case, and therefore is dominated by the Langmuir circulation mode. The effect of Coriolis acceleration in this case is mainly to shape the underlying featureless flow, and thereby to cause the rolls to have axes oriented to the right of the wind (in the Northern Hemisphere). Similarly, under typical wind conditions in thermally unstable conditions, the Langmuir circulation instability mode dominates thermal instability. Langmuir circulation requires the generation of coherent vorticity in the wind (streamwise) direction. The mechanism of the Langmuir force can be seen by a simple geometric argument. Ignore Coriolis acceleration and density stratification. This would be appropriate, for example, when considering small bodies of water, such as New York’s Lake George in which Langmuir conducted his extensive experiments, after seasonal overturning of the epilimnion. The equation for mean vorticity x is then (since r us ¼ 0) @x þ ðu þ us Þ rx ¼ x rðu þ us Þ þ vT r2 x ½7 @t The long downwind extent of windrows in these circumstances suggests the motions be idealized as independent of wind direction, and the Stokes drift is parallel to the wind. In this case, the wind supplies vorticity to the water column in the crosswind direction. Now suppose this flow is perturbed by a slight local increase in the windward component of velocity, as indicated in Figure 4. Such a perturbation can be created in a number of ways, for example, by breaking waves or by fluctuations in the applied wind stress. The disturbance generates vertical vorticity of opposite signs on either side of the velocity maximum. This vorticity is rotated by the Stokes drift to produce streamwise vorticity components as shown. The sense of the streamwise vorticity is to produce convergence near the surface toward the plane (x2 ¼ 0) through the velocity maximum and downwelling motion below this line. This could be a transient feature. However, the convergence
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LANGMUIR CIRCULATION AND INSTABILITY
x3
409
x2
u (x2, x3)
3 3
x1
Us (x3)
Figure 4 Sketch illustrating instability mechanism.
transports streamwise momentum toward the plane of symmetry, accelerating the fluid at the convergence line and amplifying the original perturbation. In fact, at the plane of symmetry, @u1 @u1 ð0; x3 Þ ¼ u3 ð0; x3 Þ @t @x3
Suppose the motion is periodic in the x2-direction, or that it decays as |x2|-N. Multiplying by o1 and integrating over the water depth and over a period in the x2-direction in the first case or all x2 in the second, implies @ @t
Z Z
1 2 o dx2 dx3 ¼ 2 1
If the streamwise velocity component decreases with depth, then at a downwelling plane, u3(0, x3)o0 and the right-hand side above is positive. The streamwise acceleration amplifies the initial perturbation, and the feedback provides a mechanism for instability. This picture is confirmed by detailed computation. Furthermore, the transport of streamwise momentum toward convergence zones provides a mechanism for the observed convergence zone jets. The assumption of streamwise invariance implies that no streamwise vorticity is produced by the Eulerian shear – with streamwise invariance of the mean flow, generation of mean streamwise vorticity requires Stokes drift. This may be seen directly from the streamwise component of [7]. If the wind stress is in the x1-direction, x3 oriented vertically upward from the mean free surface, then x2 will be crosswind in a right-handed (x1, x2, x3) coordinate system with unit vectors (e1, e2, e3). Then us ¼ Us(x3)e1, q/qx1 ¼ 0, and the streamwise component of vorticity, o1, satisfies
Z Z
@Us dx2 dx3 @x3 Z Z X 3 @o1 2 vT dx2 dx3 @xj j¼2 o1 o3
The second term on the right-hand side is the dissipation of the streamwise component of enstrophy, and the first term is its production rate. Production vanishes and streamwise vorticity decays in the absence of Stokes drift. This consequence of streamwise invariance is an exact result for the instantaneous motion; the presence of waves in the instantaneous motion introduces streamwise variations, and the Stokes drift represents the residual effect of these variations in the wave-averaged equations permitting the development of averaged streamwise vorticity. It is not possible to survey details of instability characteristics for the various cases relevant to the ocean. Qualitatively, however, it can be said that growth rates are consistent with the observed formation rates of Langmuir circulation. In the absence of Coriolis acceleration and stratification, the most 2 unstable mode is in the form of rolls parallel to the @o1 @u2 o1 @u3 o1 @Us @ o1 @ 2 o1 wind, and growth is monotonic in time, so the rolls þ þ ¼ o3 þ vT þ @t @x2 @x3 @x3 @x22 @x23 are stationary. When Coriolis acceleration is
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LANGMUIR CIRCULATION AND INSTABILITY
Acoustic backscatter Min
Field Observations New instrumentation, new measurement techniques, and improved data analysis capabilities have clarified the nature of Langmuir circulation. The new methods include greatly improved current meters and sonar methods. Sonars image microbubbles that have a virtually ubiquitous presence in the upper few meters of the ocean. These bubbles are organized by Langmuir circulation, and bubble plumes collected in, and carried down by the downwelling beneath, windows produce strong sonar returns. The intensity of sonar returns permit Langmuir circulation to be visualized in extended regions of the ocean. Doppler
W (cm s−1)
Max
100
50
0
Depth (m)
significant, the underlying featureless flow has a surface velocity to the right of the wind, and the most unstable mode takes an intermediate angle between this and the direction of the wind stress. The growth rates are slightly reduced compared to the comparable problem in the absence of Coriolis effects. The rolls are no longer stationary, but travel normal to their axes, and to the right of the wind. Only modest wind speeds are required to cause instability in these ways. The unstable character of the ocean under typical winds suggests that perturbations due to a variety of sources, such as breaking waves, grow and reach finite amplitudes. Computer simulations using eqns [4] support this picture.
Range (m)
410
8
15 17:10:00
17:27:30
17:45:00
Time (UTC) 27 October 1987 Figure 5 Simultaneous sonar images of windrows and bubble plumes. The upper picture shows bands of scatterers, and the lower picture shows bubble plumes below the bands, over a period of 35 min. Reproduced by permission of American Geophysical Union. Zedel L and Farmer DM (1991) Organized structures in subsurface bubble clouds: Langmuir circulation in the upper ocean. Journal of Geophysical Research 96(5): 8895. Copyright (1991) American Geophysical Union.
U 20
10 10
V 10
0 _ 10 _ 20
Figure 6 Data from a 50-m deep mixed layer in the Pacific. See text for for description. From Weller RA and Price JF (1988) Langmuir circulation within the oceanic mixed layer. Deep-Sea Research 35: 711–747.
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LANGMUIR CIRCULATION AND INSTABILITY
sonar can be used to measure speeds in the current system. Bubble plumes are found to extend to depths of several meters, often up to 10–12 m, below windrows. Since the micron-sized air bubbles providing the best sonar return signal go into solution at depths of this order, substantially deeper plumes do not exist. The depth of these plumes therefore provides a lower bound estimate on the strength of the downwelling motion. Figure 5 shows an example in which bubbles are organized into bands by Langmuir circulation, and plumes are formed beneath these bands. A quite remarkable set of measurements taken in 1982 is summarized in Figure 6. The data were taken by a string of current meters in a mixed layer about 50 m deep. Velocity measurements are shown near the surface and at 23-m depth, about halfway down from the surface to the base of the mixed layer. Tracks on the surface schematically indicate the
Figure 7 Sulfur powder released from an aircraft and drawn into bands of Langmuir circulation at time of (top to bottom) deployment, 1.5 min after, and 20 min after. Photograph widths are about 40 m and wind speed is 6 m s 1. From McLeish W (1968) On the mechanisms of wind-slick generation. Deep-Sea Research 5: 461–469.
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disposition of computer cards distributed on the surface to provide visual makers of windrows, which lie about 151 to the right of the wind. The presence of surface jets is indicated, with maximum values of about 20 cm s 1. Surprisingly, the downwelling speeds at 23 m are also about 20 cm s 1. Data taken at other times also produced downwelling speeds from 20 to 30 cm s 1 at mid-depth in the mixed layer. These large events occurred intermittently, with less impressive activity, on the order of half as intense, occurring at other times. During these observations, the wind speeds were on the order of 8–18 m s 1. Downwelling was observed to be as large as 3 cm s 1 under winds as light as 1.5 m s 1.
Figure 8 Computer simulation of surface tracers for a mixed layer depth of 50 m and wind speed of 9 m s 1. Top panel: Initial uniform distribution of tracers. Middle: distribution after 5 min. Bottom: distribution after 20 min, separation of windrows is 150 m. From Yang G and Leibovich S (1994) Large-eddy simulation of Langmuir circulation. Bulletin of the American Physical Society 39(9): 1885.
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The process of flotsam collection by Langmuir circulation is shown in Figure 7. Sulfur powder was spread along a line across the wind, and appears in an aerial photograph as the light blob in the upper panel immediately afterward. The second and third panels show the disposition of the dust after 1.5 min and 20 min, respectively. Surface tracers advected by the flow produced by a computer simulation of [4], [6], with a Smagorinsky turbulence subgrid scale closure model are shown in Figure 8. Note the similarity to Figure 7. Simulations carried out exactly the same way but omitting the Langmuir force fail to show these features. Langmuir circulation is now well-established as an important mechanism, affecting mixing and dispersion in the upper ocean and providing a dynamical link between the wind-wave field and the development of the mixed layer.
See also Breaking Waves and Near-Surface Turbulence. Bubbles. Upper Ocean Mixing Processes. Upper Ocean Time and Space Variability. Wind- and Buoyancy-Forced Upper Ocean.
Further Reading Langmuir I (1938) Surface motion of water induced by wind. Science 87: 119--123. Leibovich S (1983) The form and dynamics of Langmuir circulation. Annual Review of Fluid Mechanics 15: 391--427.
McLeish W (1968) On the mechanisms of wind-slick generation. Deep-Sea Research 15: 461--469. McWilliams JC, Sullivan PP, and Moerig C-H (1997) Langmuir turbulence in the ocean. Journal of Fluid Mechanics 334: 1--30. Phillips OM (1977) The Dynamics of the Upper Ocean, 2nd edn. Cambridge, UK: Cambridge University Press. Pollard RT (1977) Observations and theories of Langmuir circulations and their role in near surface mixing. In: Angel M (ed.) A Voyage of Discovery: G. Deacon 70th Anniversary Volume, Re, pp. 235--251. Oxford, UK: Pergamon. Simecek-Beatty D, Lehr WJ, Lai R, and Overstreet R (2000) Langmuir circulation and oil spill modelling. Spill Science and Technology Bulletin 6(3–4): 207--279. Skyllingstad ED and Denbo DW (1995) An ocean largeeddy simulation of Langmuir circulation and convection in the surface mixed layer. Journal of Geophysical Research 100: 8501--8521. Szeri AJ (1996) Langmuir Circulations on Rodeo Lagoon. Monthly Weather Review 124(2): 341--342. Thorpe SA (1985) Small-scale processes in the upper ocean boundary layer. Nature 318: 519--522. Thorpe SA (2004) Langmuir circulation. Annual Review of Fluid Mechanics 36: 55--79. Weller RA and Price JF (1988) Langmuir circulation within the oceanic mixed layer. Deep-Sea Research 35: 711--747. Yang G and Leibovich S (1994) Large-eddy simulation of Langmuir circulation. Bulletin of the American Physical Society 39(9): 1885. Zedel L and Farmer DM (1991) Organized structures in subsurface bubble clouds: Langmuir circulation in the upper ocean. Journal of Geophysical Research 96(5): 8895.
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LARGE MARINE ECOSYSTEMS K. Sherman, Narragansett Laboratory, Narragansett, RI, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1462–1469, & 2001, Elsevier Ltd.
Introduction Coastal waters around the margins of the ocean basins are in a degraded condition. With the exception of Antarctica, they are being degraded from habitat alteration, eutrophication, toxic pollution, aerosol contaminants, emerging diseases, and overfishing. It has also been recently argued by Pauly and his colleagues that the average levels of global primary productivity are limiting the carrying capacity of coastal ocean waters for supporting traditional fish and fisheries and that any further large-scale increases in yields from unmanaged fisheries are likely to be at the lower trophic levels in the marine food web and likely to disrupt marine ecosystem structure.
Large Marine Ecosystems Approximately 95% of the world’s annual fish catches are produced within the geographic boundaries of 50 large marine ecosystems (LMEs) (Figure 1A). The LMEs are regions of ocean space encompassing coastal areas from river basins and estuaries out to the seaward boundary of continental shelves, and the outer margins of coastal currents. They are relatively large regions, on the order of 200 000 km2 or greater, characterized by distinct bathymetry, hydrography, productivity, and trophically dependent populations. The close linkage between global ocean areas of highest primary productivity and the locations of the large marine ecosystems is shown in Figure 1B. Primary productivity at the base of marine food webs is a critical factor in the determination of fishery yields. Since the 1960s through the 1990s, significant changes have occurred within the LMEs, attributed in part to the affects of excessive fishing effort on the structure of food webs in LMEs.
Food Webs and Large Marine Ecosystems Since 1984, a series of LME conferences, workshops, and symposia have been held during the annual meeting of the American Association for the
Advancement of Science (AAAS). In the subsequent intervening 15 years, 33 case studies of LMEs were prepared, peer-reviewed, and published (see Further Reading). From the perspective of actual and potential fish yields of the LMEs an ‘ECOPATH’-type trophic model, based on the use of a static system of linear equations for different species in the food web, has been developed by Polovina, Pauly and Christensen (eqn [1]). Pi ¼ Exi þ
X
Bi ðQ=Bi Þ DCji þ Bi ðP=BÞ ðIEEi Þ ½1
Pi is the production during any normal period (usually one year) of group i; Exi represents P the exports (fishery catches and emigration) of i; i represents the summation over all predators of I; Bj and Bi are the biomasses of the predator J and group I, respectively; Q=Bj is the relative food consumption of j; DCji the fraction that i constitutes of the diet of Bi is the biomass of i and (I EEi) is the other mortality of I, that is the fraction of i’s production that is not consumed within or exported from the system under consideration. A practical consideration of food web dynamics in LMEs is the effect that changes in the structure of marine food webs could have on the long-term sustainability of fish species biomass yields.
Biomass Yields and Food Webs South China Sea Large Marine Ecosystem
An example of the use of fisheries yield data in constructing estimates of combined prey consumption by trophic levels is depicted in Figure 2 for shallow waters of the South China Sea (SCS) LME. The trophic transfers up the food web from phytoplankton to apex predators is shown in Figure 3 for open-ocean areas of the SCS. The differences in fish/ fish predation is approximately 50% of the fish production in the shallow-water subsystem and increases to 95% in the open-ocean subsystem. Application of the ECOPATH model to the SCS LME by Pauly and Christensen produced an initial outcome of an additional 5.8 Mt annually. This is a rate that is nearly double the average annual catch reported for the SCS up through 1993, indicating some flexibility for increasing catches from the ecosystem, but not fully realizing its potential because of technical difficulties in fishing methodologies.
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Arctic Ocean 19
17 20
10
46
47
1
22
18
2
21
48
9
23 44 42 41 43
45 North Pacific Ocean
3
6
5
11
35 36 37
7
8
North Atlantic Ocean
12
4 50 38
26
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25 33
27
32
16 28
31
39 13 40
15
34
Indian Ocean
29
South Pacific Ocean
30 South Atlantic Ocean
Indian Ocean
49
(A)
1. 2. 3. 4. 5. 6. 7. 8. 9. 10.
Eastern Bering Sea Gulf of Alaska California Current Gulf of California Gulf of Mexico South-east US Continental Shelf North-east US Continental Shelf Scotian Shelf Newfoundland Shelf West Greenland Shelf
11. 12. 13. 14. 15. 16. 17. 18. 19. 20.
Insular Pacific-Hawaiian Caribbean Sea Humboldt Current Patagonian Shelf Brazil Current North-east Brazil Shelf East Greenland Shelf Iceland Shelf Barents Sea Norwegian Shelf
21. 22. 23. 24. 25. 26. 27. 28. 29. 30.
North Sea Baltic Sea Celtic-Biscay Shelf Iberian Coastal Mediterranean Sea Black Sea Canary Current Gulf of Guinea Benguela Current Agulhas Current
31. 32. 33. 34. 35. 36. 37. 38. 39. 40.
Somali Coastal Current Arabian Sea Red Sea Bay of Bengal South China Sea Sulu-Celebes Seas Indonesian Seas Northern Australian Shelf Great Barrier Reef New Zealand Shelf
41. 42. 43. 44. 45. 46. 47. 48. 49. 50.
East China Sea Yellow Sea Kuroshio Current Sea of Japan Oyashio Current Sea of Okhotsk West Bering Sea Faroe Plateau Antarctic Pacific Central American Coastal
Figure 1 Boundaries of 50 large marine ecosystems (and) (B) SeaWiFS chlorophyll and outlines of LME boundaries.
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LARGE MARINE ECOSYSTEMS
Figure 2 South China Sea shallow-water food web based on the ECOPATH model. (From Pauly and Christensen (1993).)
Figure 3 South China Sea open-ocean food web. (From Pauly and Christensen (1993).)
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East China Sea Large Marine Ecosystem Evidence for the negative effects of fishing down the food chain can be found in the report by Chen and Shen for the East China Sea (ECS) LME. For a 30-year period of the early 1960s to the early 1990s, little change was reported in the productivity and community composition of the plankton at the lower end of the food chain of the ECS. However, during the same period major changes were reported for a shift in biomass yields among the ‘old traditional’ bottom species (yellow croaker) and new species dominated by shrimp, crab, and small pelagic fish species. It appears that the annual catch increase from 0.9 Mt in the 1960s to 5.8 Mt in the early 1990s exceeded the sustainable level of yield for several species. The greatest increases in biomass yield during this period has been in a category designated as ‘Other Species.’ The species in this category are near the base of the food web. They are relatively small, pelagic, and fast growing, and are not used for human consumption but are used for feeding ‘cultured fish or poultry’ (Figure 4). Collectively, the catches of ‘Other Species’ provide additional evidence of the effects of ‘fishing down the food web.’
anchovy population in the YS LME has significantly increased from an annual catch level of 1000 Mt in the 1960s to an estimated biomass of 4 Mt in the 1990s. Overfishing has led to major structural changes in the fish community of the YS LME. In the 1950s and 1960s bottom fish were the major target species in China’s fisheries. Small yellow croaker was the dominant preferred demersal market species in the late 1950s, constituting about 40% of research vessel trawl catches. By 1986, pelagic fish dominated the catches (B50%) of research vessel surveys suggesting that they may have replaced depleted demersal stocks and are effectively utilizing surplus zooplankton production no longer utilized by the depleted large pelagics and early life-history stages of depleted fish species.
Large Marine Ecosystem Regime Shifts, Food Webs, and Biomass Yields In the eastern Pacific, large-scale oceanographic regime shifts have been a major cause of changes in food web structure and biomass yields of LMEs. Gulf of Alaska Large Marine Ecosystem
Yellow Sea Large Marine Ecosystem
A projection of the Yellow Sea food web is given in Figure 5. The decline in the east Asian LMEs of demersal species and what appears to be ‘trophic-forcing’ down the food web hypothesized by Pauly and Christensen are apparent in the changes that have occurred over 30 years in the Yellow Sea LME (YS LME). The catch statistics indicate a rapid decline of most bottom fish and large pelagic fish from the YS LME from the 1960s through the early 1990s. Recent acoustic survey results indicate that the Japanese
Evidence of the food web effects from oceanographic forcing was reported for the Gulf of Alaska LME (GA LME). An increase in biomass of zooplankton, approaching a doubling level between two periods 1956–62 and 1980–89 has been linked to favorable oceanographic conditions leading to increases in primary and secondary productivity and subsequent increases in abundance levels of pelagic fish and squid in the GA LME; it is estimated by Brodeur and Ware that total salmon abundance in the GA LME was nearly doubled in the 1980s. California Current Large Marine Ecosystem
In contrast to the 1980–89 Gulf of Alaska increases in biomass of the zooplankton and fish biomass components of the GA LME, a declining level of zooplankton has been reported for the California Current LME (CC LME) of approximately 70% over a 45-year monitoring period. The cause according to Roemmich and McGowan appears to be an increase in water column stratification due to long-term warming. The clearest food web relationship reported related to the zooplankton biomass reduction was a decrease in the abundance of pelagic sea-birds. US North-east Shelf Large Marine Ecosystem Figure 4 East China Sea fisheries yield 1960s to early 1990s, showing increased annual catches of ‘Other Species’ used mostly for fish and poultry food. (From Chen and Shen (1999).)
The US North-east Shelf LME is an ecosystem with more structured coherence in the lower food web
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LARGE MARINE ECOSYSTEMS
417
Yield to Humans Apex predators
4th trophic level Carnivores II
3rd trophic level Carnivores I
2nd trophic level Herbivores Omnivores
1st trophic level Producers
{ { { {
Large pelagic fish (e.g., Spanish mackerel)
Cephalopoda (e.g., Loligo japonica)
Demersal fish (e.g., croaker and flatfish) Blue crab
Small pelagic fish (e.g., anchovy, half-fin anchovy)
Macruran shrimph (e.g., Crangon
affinis)
Zooplankton
Benthos
Phytoplankton
Figure 5 A simplified version of the Yellow Sea food web and trophic structure based on the main resources populations in 1985– 1986. (From Tang 1993).)
than in the Gulf of Alaska or California Current systems. Following a decade of overfishing beginning in the mid-1960s, the demersal fish stocks, principally haddock, cod, and yellowtail flounder, declined to historic low levels of spawning biomass. In addition, the herring and mackerel spawning stock levels were reduced in the mid-1970s. By the mid-1980s, the demersal fish biomass had declined to less than 50% of levels in the early 1960s. Following the 1975 extension of jurisdiction by the United States to 200 miles of the continental shelf, the rebuilding of the spawning stock biomass (SSB) of herring and mackerel commenced. Beginning in 1982 there was a sharp reduction in fishing effort from foreign vessels excluded from the newly designated US Exclusive Economic Zone (EEZ). Within four years the mackerel population recovered from just under 0.5 Mt to 1 Mt in 1986 and an estimated 2 Mt by 1994. Herring recovery was also initiated in the absence of any significant fishing
effort from 1982 to 1990 when increases in SSB went from less than 0.2 Mt to 1 Mt. An unprecedented 3.5 Mt level of herring SSB was reached by 1994. The NOAA-NMFS time-series of zooplankton collected from across the entire North-east Shelf ecosystem from 1977 to 1999 is indicative of an internally coherent structure of the zooplankton component of the North-east Shelf ecosystem. During the mid to late 1990s and the unprecedented abundance levels of SSB of herring and mackerel, the zooplankton component of the ecosystem showed no evidence of significant changes in biomass levels with annual values close to the long-term annual median of 30 ml/100 m3 for the North-east Shelf ecosystem. In keeping with the robust character of the zooplankton component is the initiation of spawning stock recovery subsequent to reductions in fishing effort for cod, haddock, and yellowtail flounder. Accompanying the recovery of spawning stock biomass is the production of a strong year-class of
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haddock in 1998 and a strong year-class of yellowtail flounder in 1997. The initial increases in skate and spiny dogfish populations following the declines in cod, haddock, and flounder stocks have been significantly reduced by targeted fisheries on these species. The reductions in abundance of these predators coupled with the robust character at the lower parts of the North-east Shelf food web enhance probability for recovery of the depleted cod, haddock, and yellowtail flounder stocks.
Large Marine Ecosystem Food Web Dynamics and Biomass Yields Two major sources of long-term changes in biomass near the top of the food web—fishes and pelagic birds—have been observed and reported in the literature. In the case studies of the Yellow Sea and East China Sea, the multidecadal shift in fish community structure resulting from overfishing appeared to promote the production of small pelagic fish species, indicative of ‘fishing down the food web’ as hypothesized by Pauly and Christensen, as the abundance levels of predator species decline through overfishing. For the South China Sea, estimates from a Pauly and Christensen ECOPATH model suggests that the mean annual biomass yield of fish was not fully utilized. It appears from the case study that a significant percentage of an additional 5 Mt could be fished if managed in a sustainable manner. In the eastern Pacific the results of oceanographic regime shifts had direct impact in increasing zooplankton and fish biomass in the Gulf of Alaska LME, whereas a multidecadal warming trend in the California Current LME lowered productivity at the base of the food web and resulted in a decrease in pelagic bird biomass. The importance of fish and fisheries to the structure of marine food webs is also an important cause of variability in biomass yields. A clear demonstration of this relationship is found in the application of the ECOPATH model to four continental shelf ecosystems, where it was shown that fish preying on other fish was a principal source of fish biomass loss. The level of predation ranged from 3 to 35 times the loss to commercial fisheries. Fish are keystone components of food webs in marine ecosystems. The worldwide effort to catch fish using highly effective advanced electronics to locate them, and efficient trawling, gill-netting, and longline capture methodologies, has had an impact on the structure of marine food webs. From case studies examined, evidence indicates that the fishing effort of countries bordering on LMEs has resulted in changes in the structure of marine food webs,
ranging from significant abundance shifts in the fish component of the ecosystem from overfished demersal stocks to smaller faster-growing pelagic fish and invertebrate species (herrings, anchovies, squids) as fisheries are refocused to species down the food web, predation pressure increases on the plankton component of the ecosystem. The economic benefits to be derived from the trend in focusing fisheries down the food web to lowpriced small pelagic species used, in part, for poultry, mariculture, and hog food are less than from earnings derived from higher-priced groundfish species, raising serious questions regarding objectives of ecosystem-based management integrity of ecosystems and sustainability of fishery resources. These are questions to be addressed in the new millennium with respect to the implementation of management practices. As in the case of the US North-east Shelf LME, overfished species can recover with the application of aggressive management practices, when supported with knowledge that the integrity of the lower parts of the food web remain substantially unchanged during the recovery period. However, under conditions of recent large-scale oceanographic regime shifts in the Pacific, evidence indicates that the biodiversity and biomass yields of the north-east sector of the Pacific in the Gulf of Alaska LME were significantly enhanced from increased productivity through the food web from the base to the zooplankton and on to a doubling of the fish biomass yields close to the top level of the food web. In contrast, in the California Current ecosystem the apparent heating and deepening of the thermocline effectively reduced phytoplankton and zooplankton production over a 40-year period, suggesting that in upwelling regions prediction of oceanographic events effecting food web dynamics require increased commitment to long-term monitoring and assessment practices if forecasts on effects of regime shifts on biomass yields are to be improved. If ecosystem-based management is to be effective, it will be desirable to refine ECOPATH-type models for estimating the carrying capacity of LMEs in relation to sustainability levels for fishing selected species. It was assumed in the early 1980s by Skud, based on the historic record, that herring and mackerel stocks inhabiting the US North-east Shelf ecosystem could not be supported at high biomass levels simultaneously by the carrying capacity of the ecosystem. However, subsequent events have demonstrated the carrying capacity of the ecosystem is now of sufficient robustness to support an unprecedented almost 5.5 Mt of spawning biomass of both species combined. In addition, the ecosystem in its present state apparently has the carrying capacity
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LARGE MARINE ECOSYSTEMS
to support the growing spawning biomass of recovering haddock and flounder stocks. Evidence of the production of strong year-classes for both species supported by high average levels of primary production of 350 g Cm2 y1, a robust level of zooplankton biomass, relatively high levels of epibenthic macrofauna, and apparent absence of any large-scale oceanographic regime shift suggests that integrity of the ecosystem food web will enhance the return of the fish component of the ecosystem to the more balanced demersal–pelagic community structure inhabiting the shelf prior to the massive overfishing perturbation of the 1960s to the 1980s.
Prospectus: Food Webs and Large Marine Ecosystem Management It is clear from the LME studies examined that timeseries measurements of physical oceanographic conditions that are coupled with appropriate indicators of food web integrity (e.g., phytoplankton, chlorophyll primary productivity, zooplankton, fish demography) are essential components of a marine science program designed to support the newly emergent concept of ecosystem-based management. It is important to consider the dynamic state of LMEs and their food webs in considering management protocols, recognizing that they will need to be considered from an adaptive perspective. To assist economically developing countries in taking positive steps toward achieving improved understanding of food web dynamics and their role in contributing to longer-term sustainability of fish biomass yields, reducing and controlling coastal pollution and habitat degradation, and improving oceanographic and resource forecasting systems, the Global Environment Facility (GEF) and its $2 billion trust fund has been opened to universal participation that builds on partnerships with several UN agencies (e.g., World Bank, UNDP, UNEP, UNIDO). The GEF, located within the World Bank, is an organization established to provide financial support to post-Rio Conference actions by developing nations for improving global environmental conditions in accordance with GEF operational guidelines.
See also Fish: Demersal Fish (Life Histories, Behavior, Adaptations). Dynamics of Exploited Marine Fish Populations. Ecosystem Effects of Fishing. Fish Larvae. Fisheries and Climate. Fisheries: Multispecies Dynamics. Fisheries Overview. International Organizations. Marine Fishery Resources, Global State of. Network Analysis of Food Webs.
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Pelagic Biogeography. Pelagic Fishes. Plankton. Population Dynamics Models. Primary Production Distribution. Satellite Remote Sensing: Ocean Color. Seabirds and Fisheries Interactions. Upper Ocean Time and Space Variability. Upwelling Ecosystems.
Further Reading Chen Y and Shen X (1999) Changes in the biomass of the East China Sea ecosystem. In: Sherman K and Tang Q (eds.) Large Marine Ecosystems of the Pacific Rim: Assessment, Sustainability, and Management, pp. 221--239. Malden MA: Blackwell Science. Kumpf H, Stiedinger K, and Sherman K (1999) The Gulf of Mexico Large Marine Ecosystem: Assessment, Sustainability, and Management. Malden, MA: Blackwell Science. Pauly D and Christensen V (1993) Stratified models of large marine ecosystems: a general approach and an application to the South China Sea. In: Sherman K, Alexander LM, and Gold BD (eds.) Large Marine Ecosystems: Stress, Mitigation, and Sustainability, pp. 148--174. Washington DC: AAAS. Sherman K and Alexander LM (eds.) (1986) Variability and Management of Large Marine Ecosystems, AAAS Selected Symposium 99. Boulder, CO: Westview Press. Sherman K and Alexander LM (eds.) (1989) Biomass Yields and Geography of Large Marine Ecosystems, AAAS Selected Symposium 111. Boulder, CO: Westview Press. Sherman K, Alexander LM, and Gold BD (eds.) (1990) Large Marine Ecosystems: Patterns, Processes, and Yields, AAAS Symposium. Washington, DC: AAAS. Sherman K, Alexander LM, and Gold BD (eds.) (1991) Food Chains, Yields, Models, and Management of Large Marine Ecosystems. AAAS Symposium Boulder, CO: Westview Press. Sherman K, Alexander LM, and Gold BD (eds.) (1992) Large Marine Ecosystems: Stress, Mitigation, and Sustainability. Washington, DC: AAAS. Sherman K, Jaworski NA, and Smayda TJ (eds.) (1996) The Northeast Shelf Ecosystem: Assessment, Sustainability, and Management. Cambridge, MA: Blackwell Science. Sherman K, Okemwa EN, and Ntiba MJ (eds.) (1998) Large Marine Ecosystems of the Indian Ocean: Assessment, Sustainability, and Management. Malden, MA: Blackwell Science. Sherman K and Tang Q (eds.) (1999) Large Marine Ecosystems of the Pacific Rim: Assessment, Sustainability, and Management. Malden, MA: Blackwell Science. Tang Q (1993) Effects of long-term physical and biological perturbations on the contemporary biomass yields of the Yellow Sea ecosystem. In: Sherman K, Alexander LM, and Gold BD (eds.) Large Marine Ecosystems: Stress, Mitigation, and Sustainability, Washington DC, AAAS Press, pp. 79--93.
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LARIDAE, STERNIDAE, AND RYNCHOPIDAE J. Burger, Rutgers University, Piscataway, NJ, USA M. Gochfeld, Environmental and Community Medicine, Piscataway, NJ, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1469–1480, & 2001, Elsevier Ltd.
Introduction Gulls belong to the family Laridae and terns to the family Sternidae, although many authorities treat them as subfamilies (Larinae, Sterninae) of a single family, the Laridae. Skimmers belong to the Rynchopidae. Gulls, terns, and skimmers, all members of the order Charadriiformes, are similar in many respects, but they differ significantly in many morphological, behavioral, and ecological ways. Gulls and terns are generally diurnal species that perform most of their breeding, foraging, and migrating activities during the day, while skimmers are largely nocturnal, and forage and court mostly at night. The only gulls that are primarily nocturnal during the breeding season are the swallow-tailed gull of the Galapagos Islands and the gray gull that breeds in the deserts of northern Chile. Of all seabirds, gulls are among the least specialized, and occupy a wide variety of habitats from the high Arctic and subAntarctic islands, to tropical sea coasts, and even to interior marshes and deserts. In both breeding and feeding, gulls are generalists, and their overall body shape reflects their lack of specialization to any one foraging method, food type, or nesting habitat. Gulls are highly gregarious birds that breed, roost, feed, and migrate in large colonies or flocks. Terns and skimmers are more specialized than gulls, both in their breeding habitat and in their foraging behavior, and skimmers have a highly specialized morphology and feeding behavior. While individual species of gulls, such as herring gull, may breed in many different habitats, ranging from dry land to cliffs, species of terns and skimmers breed in fewer habitats, and some are quite stereotypic in their habitat selection. Gulls feed in more different habitats on many different foods, while terns feed mainly over water by plungediving or dipping. Skimmers have one of the most unique feeding methods, skimming the water surface.
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Taxonomy The gulls are a worldwide group of about 51 currently recognized species with the main diversity occurring in both north and south temperate latitudes. Terns are also a worldwide group of about 44 species, with the main diversity occurring in tropical as well as temperate latitudes, while each of the three species of skimmers has a more limited distribution, one each in the Americas, Africa, and Asia (scientific names given in Tables 1 and 2). There is a tendency for taxonomists working at higher categories to lump genera and families together, where specialists on particular groups are more likely to emphasize differences within the group, by generic splitting – the approach followed here. Gulls
There are several natural subgroups among the gulls, most of which can be assigned either to the large white-headed or the small dark-hooded tribes. On behavioral grounds, emphasizing the commonality of display patterns, Moynihan treated all gulls in the genus Larus. Most taxonomists, however, separate some relatively unique gulls into their own genera, including the swallow-tailed gull (Creagrus), of the Galapagos, and several Arctic species, including Ross’s gull (Rhodostethia), ivory gull (Pagophila), kittiwakes (Rissa) and Sabine’s gull (Xema). Less often two south temperate species, the dolphin gull (Leucophaeus) and occasionally the Pacific gull (Gabianus) are separated as well. Terns
The main groups of terns include the black-capped terns (mostly in the genus Sterna), marsh terns (Chlidonias), noddies (Anous, Procelsterna, Gygis), and the Inca tern (Larosterna). The capped terns include small and medium sized birds in the genus Sterna and the distinctive, large, crested terns (appropriately considered a separate genus Thalasseus, although often placed into Sterna). Some unique capped terns are the gull-billed (Gelochelidon), Caspian (Hydroprogne), and large-billed (Phaetusa) terns. Sterna is a relatively homogenous assemblage. Noddies are uniformly colored, either all dark (black, brown, blue and gray noddies) or all white (white tern), while the unique Inca tern is all dark with dramatic yellow wattles at the gape.
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Table 1
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Gulls of the world with general breeding range
Gull
Scientific name
General breeding range (winter range)
Dolphin gull Pacific gull Band-tailed gull Olrog’s gull Black-tailed gull Gray gull Heermann’s gull White-eyed gull Sooty gull Mew gull Audouin’s gull Ring-billed gull California gull Great black-backed gull Kelp gull Glaucous-winged gull Western gull Yellow-footed gull Glaucous gull Iceland gull Thayer’s gull Herring gull Yellow-legged gull Armenian gull Slaty-backed gull Lesser black-backed gull Greater black-headed gull Brown-headed gull Gray-headed gull Hartlaub’s gull Silver gull Red-billed gull Black-billed gull Brown-hooded gull Black-headed gull Slender-billed gull Bonaparte’s gull Saunder’s gull Andean gull Mediterranean gull Relict gull Lava gull Laughing gull Franklin’s gull Little gull Ivory gull Ross’s gull Sabine’s gull Swallow-tailed gull Black-legged kittiwake Red-legged kittiwake
Leucophaeus scoresbii Larusa pacificus Larus belcheri Larus atlanticus Larus crassirostris Larus modestus Larus heermanni Larus leucophthalmus Larus hemprichii Larus canus Larus audouinii Larus delawarensis Larus californicus Larus marinus Larus dominicanus Larus glaucescens Larus occidentalis Larus livens Larus hyperboreus Larus glaucoides Larus thayerib Larus argentatus Larus cachinnans Larus armenicusb Larus schistisagus Larus fuscus Larus icthyaetus Larus brunnicephalus Larus cirrocephalus Larus hartlaubii Larus novaehollandiae Larus scopulinus Larus bulleri Larus maculipennis Larus ridibundus Larus genei Larus philadelphia Larus saundersi Larus serranus Larus melanocephalus Larus relictus Larus fuliginosus Larus atricilla Larus pipixcan Larus minutas Pagophila eburnea Rhodostethia rosea Xema sabini Creagrus furcatus Rissa tridactyla Rissa brevirostris
Southern South America Southern Australia Pacific Coast of South America Uruguay to Argentina Siberia to Japan (east China) Western South America (coastal) West Mexico (California) Red Sea Middle East (East Africa and Pakistan) Eurasia and western North America (Africa, east Asia) Western Mediterranean and West Africa Central and eastern North America (to northern South America) Central North America (Pacific Coast to west Mexico) Holarctic (south-eastern US, western Europe) Southern South America, Australia, Africa Eastern Asia to western North America Pacific Coast of North America Gulf of California Holarctic Eastern Canadian Arctic and Greenland (North America and Europe) Arctic Canada (Pacific Coast of North America) Holarctic (to northern South America and South-east Asia) Eurasia (North Africa, Middle East to Western India) Armenia, Turkey (Middle East) Siberia to Japan (south to Taiwan, few in Alaska) Eurasia (Africa, Arabia, south-western Asia) Central Asia (southern and south-east Asia) Central Asia (southern and south-east Asia) Southern South America, Africa South Africa Australia, New Caledonia New Zealand New Zealand Southern South America (northward along coasts) Eurasia (Africa, southern Asia) few in North America Southern Europe to western Asia (Middle East) Alaska–Canada (North American coasts to Caribbean) North-eastern China (China to Japan) Andes (Pacific Coast of South America) Southern Europe (north Africa) Central Asia (unknown) Galapagos endemic Eastern US, Caribbean to South America Central North America (South America) Eurasia (Europe, North Africa) Holarctic Siberia to Alaska Holarctic Galapagos and Malpelo I (Colombia) (South America) Holarctic Bering Sea endemic
a b
Often placed in the monotypic genus Gabianus. The status of this as a species is in question.
Skimmers
The three species of skimmers are closely related and form a superspecies. The Indian skimmer breeds
in southern Asia from Pakistan to Cambodia; the African skimmer breeds along rivers in Africa, and the black skimmers breeds in North and South America.
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Table 2
Terns and skimmers of the world with general breeding range
Tern/skimmer
Scientific name
General breeding range (winter range)
Gull-billed tern Caspian tern Elegant tern Lesser crested tern Sandwich tern Chinese crested tern Royal tern Greater crested tern Large-billed tern River tern Roseate tern Black-naped tern White-fronted tern South American tern Arctic tern Antarctic tern Kerguelen tern Forster’s tern Trudeau’s tern Little tern Saunder’s tern Least tern Yellow-billed tern Peruvian tern Fairy tern Damara tern White-cheeked tern Black-bellied tern Aleutian tern Gray-backed tern Bridled tern Sooty tern Black-fronted tern Whiskered tern White-winged tern Black tern Brown noddy Black noddy Lesser noddy Blue noddy Gray noddy White tern Inca tern Black skimmer Africa skimmer Indian skimmer
Gelochelidona nilotica Hydroprognea caspia Thalasseusa elegans Thalasseus bengalensis Thalasseus sandvicensis Thalasseus bernsteini Thalasseus maximus Thalasseus bergii Phaetusa simplex Sterna aurantia Sterna dougallii Sterna sumatrana Sterna striata Sterna hirundinacea Sterna paradisaea Sterna vittata Sterna virgta Sterna forsteri Sterna trudeaui Sterna albifrons Sterna saundersi Sterna antillarum Sterna superciliaris Sterna lorata Sterna nereis Sterna balaenarum Sterna repressa Sterna acuticauda Sterna aleutica Sterna lunata Sterna anaethetus Sterna fuscata Sterna albistriata Chlidonias hybridus Chlidonias leucopterus Chlidonias niger Anous stolidus Anous minutus Anous tenuirostris Procelsterna caeruleab Procelsterna albivittab Gygis alba Larosterna inca Rynchops niger Rynchops flavirostris Rynchops albicollis
Old and New World Old and New World Pacific Coast of North America (to South America) Old World tropics American tropics and Europe (Africa) China: Taiwan; nearly extinct (Indonesia) American tropics and West Africa Old World tropics South American rivers (coastal South America) India and south-east Asia Pan-tropical Old World tropics New Zealand (south-east Australia) South American coasts Holarctic (subAntarctic) Subantarctic (southern Africa) Crozets, Prince Edward, Marion, Kerguelen North America (Central America) Southern South America Eurasia, Africa, Australasia Indian Ocean (Africa) North America (northern South America) South American rivers (coastal South America) Pacific Coast of South America Western and southern Australia South Africa (to West Africa) Indian Ocean and Middle East India and south-east Asia Siberia and Alaska Central and western Pacific Pan-tropical Pan-tropical New Zealand rivers (New Zealand coast) Eurasia, Africa, Australasia Eurasia (Africa, Australasia) North America and Europe (South America, Africa) Pan-tropical West Indies, Australia, south-west Pacific Indian Ocean to western Australia Tropical western Pacific Temperate Pacific Pan-tropical Pacific coast of South America North American coasts and South America inland West, central, and East Africa (Egypt) North Indian subcontinent to Cambodia
a b
These genera are often placed in the genus Sterna. These two species have until recently been considered a single species.
Physical Appearance In all species of these families, males and females are indistinguishable on the basis of plumage. Moreover in gulls and terns there is very little sexual size dimorphism, while skimmers are among the most sexually dimorphic in size of any bird. The skimmers are heavy-bodied, with very long narrow wings, and large, narrowly compressed or knife-like bills for
skimming the water. Adult gulls, terns, and skimmers of most species have the coloration of plunge-diving sea birds. They are generally white below, which is believed to serve as camouflage against the pale sky, reducing their conspicuousness to their underwater prey. Young birds are generally spotted or blotched or streaked, affording camouflage on the various substrates they occupy, particularly during the
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Figure 1 Examples of gulls, terns and skimmers. (1) common tern (Sterna hirundo) Adult, breeding plumage. Length: 36 cm; wingspan: 80 cm; approximate body mass 120 g. Range: Widely distributed in Northern Hemisphere during breeding season, and coastal oceans of the world during the non-breeding season. (2) Sooty tern (Sterna fuscata). Other names: Wideawake. (a) Adult; (b) juvenile. Length: 44 cm; wingspan: 90 cm; approximate body mass: 195 g. Range: Tropical Oceans. (3) Black skimmer (Rynchops niger). Adult, breeding. Length: 45 cm; wingspan: 117 cm; Range: Coastal waters and some rivers of North and South America. (4) Herring gull (Larus argentatus). Adult, breeding plumage. Length: 61 cm; wingspan: 139 cm; approximate body mass: 1160 g. Range: Widespread in costal waters of North Pacific and North Atlantic Oceans. Also along coasts of Mediterranean Sea and Indian Ocean. (5) Great black-backed gull (Larus marinus). Adult breeding plumage. Length: 75 cm; wingspan: 160 cm; approximate body mass: 1680 g. Range: North Atlantic Ocean. (6) Kittiwake (Rissa tridactyla). Other name: black-legged kittiwake. Adult breeding plumage. Length: 41 cm; wingspan: 91 cm; approximate body mass: 385 g. Range: North Pacific, North Atlantic and Arctic Oceans. Illustrations from Harrison P (1985) Seabirds, an identification guide. Revised edition. (Boston, Massachusetts: Houghton Mifflin).
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critical pre-fledging period when they rely on cryptic coloration to avoid predation (Figure 1). Gulls
As a group gulls are generally heavy-bodied, longwinged birds with intermediate length necks and tarsi, webbed feet, and heavy, slightly-hooked bills. Gulls range in weight from as little as 100 g for a little gull to 2000 g for some great black-backed gulls. Gulls exhibit no sexual dimorphism in plumage patterns, but females are slightly smaller than males and in some species show more slender necks and beaks. All species have 12 rectrices and the tails are rounded in all but a few species (Sabine’s gull, swallow-tailed gull). Gulls are generally white-bodied, with a darker mantle varying from pale silvery-gray to black (except for the all-white ivory gull). Several of the smaller species exhibit a pale pink or cream-colored bloom on their breast early in the breeding season. In the Ross’s gull (not roseate tern) this is very pronounced and persistent. Gulls have either a dark hood (nearly obscure in some species), dark mask, or all white head during the breeding season. Generally the larger gulls are white-headed, although the great black-headed gull, one of the largest species, is an exception. The smaller gulls are generally either dark-masked or dark-headed. In almost all species the wingtips are black, the melanin pigment offering resistance to wear. Most species have a complex pattern of white ‘windows’ on the black outer primaries which may differ among closely related species. Several Arctic species (glaucous, Iceland, ivory gulls), have white wing tips. Some species of gulls reach adult plumage in 2 years (the smaller hooded and masked gulls), while the larger species have 3–4 distinct plumage year classes. Some do not reach fully adult plumage until their fifth year. Most gulls molt their flight feathers twice a year, and their body feathers once a year. Franklin’s gull, which migrates farther than any other gull from its breeding range in the prairies of North America to South America, undergoes two complete molts each year. However, Sabine’s gull migrates almost as far, but has only one complete molt each year. Terns
Terns have narrower, more elongated bodies than gulls and proportionately longer, more slender and pointed wings. Their beaks are generally slender and sharply pointed. Most species of terns forage at least occasionally, by plunge-diving for fish, and accordingly their body is stream-lined. Most terns are white
below and gray above, with a black crown in nuptial plumage, although a few species are all dark or all white. The roseate tern has a pale pinkish ‘bloom’ on the breast early in the breeding season. This is quickly lost by wear. There are no sexual differences in plumage patterns. The plumage patterns of the gull-billed tern, crested terns, typical black-capped, and large-billed tern are similar: mainly white below, gray above, with a black cap during the breeding season. During the nonbreeding season most species lose part or all of the black crown. Terns have one complete molt and one body molt each year. Skimmers
Skimmers are the size of the large terns, and are black above and pure white below, although there is a pale cream-colored tinge particularly on the flanks, early in the nesting season. The Indian skimmer has a broad white collar and the other two species gain such a collar in their post-breeding molt. The beak is bright reddish-orange with a yellow tip in the African and Indian skimmers and is red at the base with the distal half black in the black skimmer. The knifelike bill is adapted for slicing through the water surface when the birds are skimming to catch fish. A most unusual feature of skimmers is that the upper mandible is shorter than the lower. Skimmers have two molts a year, including a complete molt and a body molt.
Habitat Gulls and terns are quite variable in their habitat preference, although almost all species feed and nest in association with water. Gulls forage in a wider range of terrestrial and aquatic habitats, while terns and skimmers forage almost exclusively over water. Exceptions include the gull-billed tern which feeds extensively on insects obtained by plunging to the surface or hover-dipping over grasslands. Gulls
Gulls use a wide variety of habitats; for any single species the range may be narrow or broad. In the breeding season, colonies of nesting gulls can be found in coastal and estuarine habitats, as well as inland. Most species of gulls nest along the continental coasts or on large inland lakes, showing a strong preference for islands or inaccessible sites. A few (brown-hooded, Franklin’s, relict) nest mainly on inland lakes or marshes, while two, the lava and swallow-tailed gulls nest on remote oceanic islands, the Galapagos. Franklin’s gulls breed on inland
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LARIDAE, STERNIDAE, AND RYNCHOPIDAE
freshwater marshes but winter on South American sea coasts. Gulls occupy a wide variety of nesting habitats, including sandy or rocky islets or beaches, with or without vegetation, marshes, riverine or lake sand bars, wind-swept sand dunes and cliffs, trees, and even buildings. The unique gray gull breeds in the barren, montane deserts of Chile, flying each day over the Andes to the Pacific Coast to obtain food for their young. Gull colonies are generally located in habitats removed from mammalian predators, but some gulls will breed on terrestrial habitats where both avian and mammalian predators are threats. During migration gulls fly to coastal and estuarine habitats, and in winter, they generally remain along coasts or on large lakes. Although they can be found out to the continental shelf, most are not truly pelagic. During the winter, their distribution is a function of food availability. Since they eat a wide variety of foods, they can be found in nearly all aquatic habitats. Outside the breeding season, gulls are found at virtually all latitudes where open water is available. Terns
Terns occur throughout the world, and breed on all continents, including Antarctica (Antarctic tern). Although individual tern species are less variable in their breeding habitats than gulls, overall, terns occupy a wide range of breeding habitats, including inland and coastal marshes, islands in rivers, lakes and estuaries, sandy or rocky beaches, cliffs, and oceanic islands. Since most terns nest on the ground, they seek remote or inaccessible places as an antipredator strategy against mammalian predators. Several of the noddy species nest on trees. During the nonbreeding season, most species of terns migrate to coastal estuaries and the open ocean, although some never leave their inland marshes and rivers. Some species are truly pelagic during most of the nonbreeding season. Sooty tern, the most oceanic tern species, maintains a pelagic existence from the time of fledging until returning to land to breed for the first time at several years of age. Arctic terns wintering in Antarctica often roam through the loose pack ice along the edge of the ice pack. Skimmers
The skimmers in South America, Africa, and Asia are largely restricted to nesting on riverine sand bars and islands in lakes. The black skimmers of North America are mainly coastal, nesting on beaches and islands, although small numbers breed inland. Except in North America, the nesting of skimmers is
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influenced by rains, for they must wait until water levels fall and their islands are exposed. Skimmers are the least pelagic members of this family, and rarely are observed far from land.
Breeding Ecology and Behavior Gulls, terns, and skimmers normally breed in colonies, either in monospecific colonies, or in monospecific groups within colonies that include other species. Terns are the most gregarious and they generally breed, forage, and migrate in flocks that can range from a few individuals to many thousands or even millions (sooty tern). Most gulls and terns in temperate zones breed at the same time of year, breed once a year, and breed every year. While North American skimmers do likewise, skimmers that are dependent on the formation of sand bars and sandy islands in rivers must wait until such sand bars are exposed. Some tropical terns have shorter breeding seasons, and can breed every 8 months. Territory size generally increases with body size for gulls, and decreases with body size for terns (Figures 2 and 3). Based on an examination of 49 species of gulls and terns nesting in aquatic and terrestrial habitats, variability in density was accounted for by colony size, body length, and wingspan, and variability in nearest neighbor distance was accounted for by body size. Gulls and terns of several species, such as common tern, can be very aggressive at mobbing potential predators, including human intruders into their colony. Such mobbing behavior can deter avian predators, but is less effective in discouraging mammalian predators. Small species often nest in colonies with larger species to take advantage of the larger size and antipredator behavior of these species, and skimmers often nest in colonies with other species to take advantage of their mobbing behavior, because they do not normally do so. Both members of the pair engage in territory defense, incubating, and protection and feeding of the young. The eggs of most gulls and terns are brown with dark splotches, while the base color of skimmer eggs is white. Clutch size in most gulls and terns is two or three, while in skimmers it is more variable, ranging from two to four or occasionally five. Since incubation begins after laying the second egg in skimmers, there is asynchronous hatching, and the young can be very variable in size. The first two chicks may hatch only a day apart, while the fourth may not hatch until 5 days later. Pairs that lose eggs or chicks may initiate a repeat nesting attempt. Following the breeding season, young skimmers and
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7000 6500 6000 5500
50
Gulls
5000
40
4500 All
4000 3500 3000 2500 2000 1500 1000
20
500 0 4500 4000 _1
10
Number of pairs ha
Nearest neighbor distance (m)
30
Terns
Gulls
3500 3000 2500 2000 1500 1000 500 0
10
9000 8000 7000
Terns
6000
30
0 0 0 5 _ 35 _ 4 _ 4 _ 5 _ 55 _ 6 _ 65 6 6 6 1 1 0 4 3 5 4 5 61 3 Body length (cm)
65
Figure 2 Relationship between body size and territory size (nearest neighbor distance) showing that smaller gulls and larger terns tend to nest more densely.
5000 4000 3000 2000 1000 0 30
gulls may remain with their parents for a few days or weeks while they improve foraging skills, however, young terns remain with their parents for weeks or months, perfecting the difficult task of plunge-diving. Some terns migrate with their parents and remain with them much of the winter. In a normal breeding season the phenology includes (1) arrival in the colony vicinity a few weeks before occupation, (2) colony occupation for a few weeks before initiation of nest-building and egglaying, (3) territorial defense and courtship during this period, (4) courtship feeding of the female by the male and selection of a nest site and nest construction, (5) a copulation period of a few days to 2 weeks, (6) an egg-laying period of 1–3 weeks, (7) an incubation period of 20–30 days resulting in incubation activities in the colony over about 6 weeks,
0 0 0 _ 55 _ 6 _ 65 _ 35 _ 4 _ 45 _ 5 46 36 56 51 61 30 41 Body length (cm)
65
Figure 3 Relationship between body size and colony size showing that smaller gulls tend to nest in larger colonies while smaller terns tend to nest in smaller colonies.
and (8) a brooding or pre-fledging phase of 4–6 weeks. These time periods are slightly less for the small gulls, and slightly more for the large gulls. Environmental constraints such as low food supply or bad weather impose synchrony on the breeding cycle. Gulls
Most species of gulls nest in a wide range of habitats, although some species are specialists, nesting only in
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LARIDAE, STERNIDAE, AND RYNCHOPIDAE
one habitat (Franklin’s and brown-hooded gulls nest only in marshes, kittiwakes nest on cliff ledges or buildings, and Bonaparte’s gulls normally nest in trees). Most gulls breed in colonies of few to several hundred pairs, although some species will occasionally breed solitarily (herring gull, Bonaparte’s gull), and others will breed in colonies of several thousand pairs (Figure 3). The small gulls generally nest in more dense colonies than do the larger gulls. The large gulls nest in more open colonies, which is partly the result of cannibalism. Species nesting in or adjacent to gull colonies include: grebes, ducks, cormorants, herons and egrets, gannets, alcids, and even penguins. Some of these apparently do so because they derive early warning and antipredator advantages from the gulls, although they also risk predation by the gulls. The nesting period usually lasts from 3 months for the smaller gulls, and up to 5 months for the larger species. A prominent exception to the annual cycle is the swallow-tailed gull of the Galapagos which breeds in all months of the year, and even within a colony there is breeding asynchrony among pairs. The age of first breeding in gulls varies from 2 (for some smaller species) to 5 years of age (for some large species). Females are likely to start breeding at a younger age than males. Factors that influence age of breeding within a species include availability of mates, nesting sites, food, and weather conditions. Reproductive success usually increases over the first few years as birds gain breeding experience; success remains relatively high for a number of years, and decreases only in very old gulls. A typical cycle is exemplified by the widespread herring gull. Early in the breeding cycle, gulls begin to gather on ‘clubs’ on their nesting colonies, where they solicit for mates or re-establish pair bonds. Gradually the gulls spread out over the nesting colony to delineate and defend nesting territories. Unmated males defend territories while courting potential mates; mated pairs defend the territory together and engage in pair-bond maintenance. Once pairs have formed, display frequency increases, and pairs may spend hours every day engaging in ‘mew calls’, mutual ‘long-calling’, and ‘head-tossing’. Within a few days, the pair begins to select a nest site, and each member shows the other its choice for a nest site with a ‘choking’ display. Nest building is an integral part of pair formation and pair maintenance for some species, and is less important for others, depending upon physical constraints. All gulls construct a nest, and use vegetation to form a nest cup. In most species males courtship-feed their mates in the pre-egg-laying period, although the level
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of courtship feeding is less than it is for terns. Copulation is always accompanied by a staccato, repetitive ‘copulatory call’ that is accompanied by ‘wing-flagging’ which is visible far across the colony. This leads to contagious copulation by neighboring pairs, and is believed to lead to greater breeding synchrony. Most gulls have a modal clutch size of three eggs, although some species such as Hartlaub’s, blackbilled, silver gulls, and kittiwakes usually lay only two eggs, and swallow-tailed gulls only one. Both members of the pair engage in incubation, which lasts an average of 24–26 days. Incubation bouts usually range from 1 to 4 h during daylight, with one member of the pair remaining on the nest at night. The first two chicks usually hatch within a few hours of one another, but the third chick may hatch a day or two later, giving it a distinct disadvantage when competing with siblings for food. The cryptic downy chicks are usually a pale gray or pale tan with dark splotches. The chicks are brooded until they are 1–2 weeks old, and usually are guarded until they fledge. Both parents feed chicks, although during the first week the male performs more provisioning, while the female remains on the territory and broods the chicks. After the first week or two, both members of the pair take turns with foraging trips and guarding the chicks. Unless disturbed by storms, floods or humans, they use the nest throughout the brooding period. The pre-fledging period in gulls ranges from 4 to 7 weeks, depending on the size of the species. The full extent of post-fledging parental care has not been established for gulls, but for some it clearly extends for a few weeks post-fledging. In all species of gulls studied, pair-bonds are monogamous, with a relatively high degree of mate fidelity from year to year. Pairing for life is characteristic of gulls that have been studied. However, divorce occurs when pairs are unable to work out incubation and brooding activities, or when they are unsuccessful at raising offspring. Pairs that remain together lay earlier, and raise more young than pairs that have found new mates. Study of the mating systems of gulls has been stimulated because of the relatively recent reports of female–female pairs in ring-billed, herring, western, red-billed, and California Gulls. Female–female pairs constitute at least 10% of the pairs of western gulls in some colonies and 6% of the pairs in red-billed gulls. Female–female pairs apparently result from a shortage of breeding males, which may be a consequence of differential survival rates. Laboratory experimentation with western gull eggs indicated DDT-induced feminization of genetically male gull embryos. The levels of DDT that caused feminization
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were of the same magnitude as those found in seabird eggs in Southern California in the early 1970s. Developmental feminization of males may be associated with inability to breed as adults, and may explain the highly skewed sex ratios and reduced number of male gulls in some populations. Some promiscuity occurs in gulls, with both sexes copulating with birds other than their mates. Males and females invest approximately equivalent effort, averaged over the entire season, although males often provide more of the initial territory defense, perform all of the courtship feeding, and may perform the bulk of territorial defense, as occurs in western gulls and ring-billed gulls. Females sometimes perform more of the incubation duties and chick defense than males. Moreover, in territorial clashes, males tend to defend against males, and females defend against females when both are present on the territory. Terns
Although terns have a worldwide distribution, some individual species have a very restricted range. Terns breed in colonies that may be used for several years or decades, and are abandoned when the breeding site becomes unsuitable because of increased predation, human disturbance, habitat changes, or weather stresses such as severe storms or floods. Philopatry varies as a function of habitat stability, and species that nest on remote oceanic islands or on cliffs occupy the same colony site for years. Changing habitats, such as freshwater marshes, riverine sand bars, and coastal sandspits, may be used for only a few years or even only 1 year. In some species individual terns are known to return to their previous site, assess its suitability, and then either occupy the site or move elsewhere. The breeding season may last only 2 months for Arctic species, 3–4 months for temperate species, and 3–5 months for tropical species. Terns usually spend only 2–3 weeks in the vicinity of the colony before they begin to settle on the colony for another breeding attempt. Temperate species of terns typically leave the breeding grounds with their chicks, and disperse before migrating to wintering grounds. The chicks may remain with their parents for many weeks or months, slowly decreasing their dependence on their parents for food. Post-fledging parental care has been documented only in those species that migrate and overwinter near land where parents can land to feed their chicks. Nesting colonies range in size from a few widely dispersed pairs (e.g. Damara and Inca terns) to colonies of a million or more (sooty terns).
Intermediate-sized terns, such as the common tern, usually nest in colonies of tens to hundreds of pairs (with a few colonies exceeding a thousand). The large terns usually nest in colonies on the order hundreds of pairs (Figure 3). All species of terns are territorial, although the size of the territory varies from 40 cm to a few meters. Unlike gulls, the inter-nest distance decreases as the size of the tern increases (Figure 2). That is, the larger species (royal, Caspian, crested, and sandwich terns) nest in very dense colonies where an incubating bird can almost reach its neighbor, the intermediate-sized terns (common, roseate, Arctic, Forster’s, noddies) nest in fairly dense colonies where inter-nest distances may range from 1 to 3 m, and the small terns often nest in fairly dispersed colonies where internest distances may be 5 m or more. The larger species rely on their dense nesting pattern to prevent any ground or aerial predators from getting within the colony, while the smaller species rely on dispersion and cryptic coloration to reduce predation, and they rely on antipredator behaviors such as mobbing and aerial attacks to repel predators. Although some species of terns nest in monospecific colonies, many species nest with other species of terns, gulls, skimmers, boobies, alcids, and ducks, and with sea turtles. Some terns (Forster’s black) choose to nest with other species, and they select the colony site after the other species is already nesting. These terns are smaller than the gulls, and derive antipredator protection from the larger gulls that mob and attack aerial predators such as hawks. Terns also nest in colonies that are adjacent to colonies of other species, such as albatrosses, boobies, and cormorants. The normal breeding phenology is to select territories, defend territories, solicit mates or reestablish pair-bonds with former mates, engage in courtship and courtship feeding for a few days or weeks, to have an egg-laying period of a few days to 2 or 3 weeks, to have an incubation period of 22–26 days, and to have a brooding and pre-fledging phase of about 4 weeks. After fledging birds may linger at the colony or depart, resulting in a colony breeding phase of up to 6–10 weeks. Smaller terns require somewhat shorter and larger terns longer, breeding times. Within a colony breeding is generally much more synchronous among Arctic species and can be much less synchronous among tropical species. The former are constrained by short breeding season and the latter by less abundant food supplies. During courtship both mated and unmated birds engage in ‘fish flights’, which begin when a mated male brings back a fish to a female. After he lands on the territory, both take off and fly high in the air,
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LARIDAE, STERNIDAE, AND RYNCHOPIDAE
often joined by one or two other birds. They glide and circle on stiffly bowed wings, uttering unique flight calls. Once pairs have formed, courtship displays continue, with the addition of courtship feeding. During this period the female often remains on the territory, defending it against intruders, while the male spends all day bringing back fish to courtship feed her. This is said to offer the female the opportunity to assess the prospective paternal qualities of her mate. Nest site selection occurs when both members of the pair agree on a given location. Most terns do not construct a nest, but merely make a scrape in the sand or roll a few pebbles or shells around a scrape, or find a suitable cup-shaped place in the coral or rock for their eggs. Terns nesting in marshes (for example Forster’s terns and black terns), construct a nest of vegetation, on which eggs can float up if flooding occurs. Tree-nesting species such as some noddies must select a branch or shelf where they can build a nest of twigs and feathers cemented by their excreta. Cliff or ledge-nesting species such as brown noddies often move the small coral rubble to the edge of the cliff, providing a smooth surface for the eggs and chicks, and a slight lip to prevent the egg from rolling off. The white tern makes no nest at all, but places its egg on a branch, or ledge, or artificial object, where it is hardly protected from falling. Clutch size in terns varies from one egg (royal, elegant, white-fronted, black-naped, sooty, noddies, white tern), to two eggs (Caspian, roseate, Arctic, sandwich, least), to three eggs (gull-billed, common). For species that normally lay two or three eggs, clutch size is dependent on food supply. In low food years, terns can reduce the average clutch size, delay breeding, or forego breeding. Members of pairs share incubation duties and chick care, although early in the brooding phase males may bring back more food and females may incubate them more often. The birds leave their overnight roost or breeding colony to search for food at dawn. These feeding flights may involve large flocks (particularly of nonbreeding birds), or prolonged streams of individuals and small groups. During the breeding season, terns that are not incubating often rest on their territory with their mates. Terns are often more active in the early morning and late afternoon. The daily activity patterns of coastal-nesting species of terns are often influenced by tidal cycles. Skimmers
The breeding behavior of skimmers is similar, except for the following: (1) they shift colony sites more often because of the changing nature of their nesting
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habitat, (2) colonies may consist of only a few pairs, (3) males defend the territory more than females, (4) a male will usually bring back a fish prior to copulation, pass it to his mate, and she may hold it while they copulate, (5) clutch size is more variable, and may range from two to five, depending upon food supply, (6) since clutch size is larger than in gulls and terns, and the parents may start incubating almost immediately, there can be greater size disparity between the offspring, and a pair rarely raises more than two young, even when food is not especially scarce, (7) males provision the chicks more often during the first 2 weeks of their lives, and (8) if disturbed, they move the chicks soon after hatching. Once chicks are more than a week old, the parents may not brood them, and they continue to be vulnerable to prolonged chilling rains, which can produce high mortality among chicks in the 1–2 week age range. At the same time the parents may be brooding the much smaller, last-hatched chicks, which may, under these circumstances, survive.
Foraging Gulls have the most diverse foraging behavior and feed on the greatest variety of foods, while terns primarily plunge-dive or hover-dip for fish, and skimmers skim the surface of the water for fish. Both terns and skimmers are limited to feeding over water, while gulls feed on land, along the shore, and over water. Both gulls and terns engage in piracy, although terns generally pirate from conspecifics or other terns, while gulls pirate from a range of species. Foraging efficiency increases with age in gulls, terns, and skimmers. Gulls forage in a variety of natural habitats including the open ocean, the surf zone, intertidal mudflats, rivers and rivermouths, rocks and jetties, estuaries, bays, lakes, reservoirs and rivers, wet meadows and farm fields, sewage outfalls, refuse dumps, and even in the air. Gulls are an integral part of coastal and estuarine habitats, and many species feed along the shore on a variety of fish and invertebrates. Gulls are particularly characteristic of the intertidal zones. They also feed in a variety of man-influenced situations, including on landfills, behind ploughs or boats, and by pan-handling from people at fast-food places or along the shore. They forage using a wide range of techniques, including walking on the ground, swimming in the water and dipping for food, and plunge-diving. They also drop invertebrates from some distance to crack open their shells, a behavior shared only with Corvids. In some species, individuals have specialized diets or foraging
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Table 3
Conservation status of selected gulls, terns, and skimmersa
Species
Status
Pacific gull Olrog’s gull
Near-threatened. Nowhere common; is declining probably due to kelp gull pressure. Vulnerable, and probably endangered. Only 1200 pairs known in six colonies, none secure. Competition from increased fishing; vulnerable to oil exploration. At least vulnerable with only eight colonies, and most of its large population in one colony. Vulnerable. Poorly known, with fewer than 7500 pairs. Vulnerable to egging, oiling, exploitation, tourism. Formerly endangered, but increasing under intense protection. Still fewer than 20 000 pairs, but now in about 30 colonies. Fishing competition potentially severe. Vulnerable due to very low numbers and small range. Probably o5000 pairs in less than a dozen colonies. Vulnerable. Low numbers and very small range in Arnebua to north-western Iran. Formerly considered a race of herring gull. Probably vulnerable. About 30 known breeding colonies, and over 25% of population in one colony (albeit protected). Endangered. Fewer than 2500 pairs (probably o1500 pairs). All seven known colony sites are being developed or slated for development. Intense management and protection urgently required. Probably endangered. Small range, few colonies, small population. Fewer than 2000 pairs and declining. Vulnerable and probably threatened. The world’s rarest gull. Although there are fewer than 400 pairs, most of the habitat is currently well-protected. Increased tourism and fishing jeopardize its survival. Vulnerable. A Bering Sea endemic that is declining. Only seven known colony sites. Although population is still large, it has declined precipitously through the 1990s. Critically endangered, breeding area unknown, often listed as extinct, but a few recent sight records on wintering grounds. Status poorly known, vulnerable Not globally threatened (c. 50 000 pairs), but North American and European populations are endangered, Caribbean population is threatened. Vulnerable and probably threatened. No truly secure colony. Total population o2500 pairs, mostly in Kerguelen where cat predation is a continuing problem. Status unknown, not known to be common anywhere. No secure colonies documented. Not globally threatened, but all races vulnerable, and California race is endangered. Status uncertain, a vulnerable species. World population probably around 5000 pairs. Vulnerable. About 5000 pairs in Australia. New Zealand population is endangered with fewer than 10 pairs. Near-threatened with fewer than 10 000 pairs. Breeding areas are neither protected nor heavily disturbed at this time. Vulnerable. Fewer than 10 000 pairs and continually declining. Vulnerable. Fewer than 10 000 pairs, mostly in small colonies on braided rivers being invaded by lupine. Not globally threatened, but status poorly known. Only a few colonies known. Not globally threatened, but eastern colonies on Desadverturas and Easter Island are at least threatened if not endangered. Not globally threatened, but fewer than 10 000 pairs, breeding in many very small colonies. Vulnerable. Declining, but poorly documented over most of its range.
Heermann’s gull White-eyed gull Audouin’s gull Yellow-footed gull Armenian gull Hartlaub’s gull Saunder’s gull Relict gull Lava gull Red-legged kittiwake Chinese crested tern River tern Roseate tern Kerguelen tern Saunder’s tern Least tern Peruvian tern Fairy tern Damara tern Black-bellied tern Black-fronted tern Lesser noddy Gray noddy African skimmer Indian skimmer a
Official global or regional status is given in boldface. The list provides both official and unofficial status of tern and gull species that are in trouble. For many of the remaining species that are globally secure, there may be many colonies, populations, or regions where it is declining or in danger of disappearance. The hallmark of security is having many colonies, including those in protected areas. The consequence of having many colonies is that not all can be protected.
techniques when compared with their populations as a whole.
Conservation Of the 99 species of gulls, terns and skimmers, world population estimates range from a few hundred pairs (lava gull) to several million pairs (herring gull) and several tens of millions (sooty tern). Saunder’s gull and Chinese crested tern are endangered. Ten species are considered ‘threatened’, and one skimmer is considered ‘vulnerable’ (see Table 3).
Although egging, hunting, and exploitation for the millinery trade resulted in sharp declines for many species in the last two centuries, human persecution has ceased in many (but not all) parts of the world. ‘Egging’, the collection of bird eggs for food, continues to be a major threat in tropical regions, (e.g. roseate terns in the Caribbean). Current threats to species in this family include habitat loss, habitat degradation, increased predation (often caused by predators, such as cats, introduced and maintained by humans), and overfishing by humans that reduces food supplies. Populations can also be threatened by pollution, particularly oil spills that occur near
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nesting colonies or in favorite foraging grounds. In the 1960s and 1970s eggshell thinning due to DDT was a problem, and in the 1980s in the Great Lakes of North America, organochlorides contributed to decreased hatching rates, lowered parental attendance, and lower reproductive success. Gulls, terns, and skimmers suffer heavy and increased predation from predators introduced by man, including cats, dogs, and by predators that have profited by man’s activities, such as foxes. In addition, many species of terns and smaller gulls have been negatively impacted by competition with larger gulls. In many temperate regions, the large whiteheaded gulls (e.g. herring gulls, ring-billed gulls) expanded their numbers and ranges dramatically during the twentieth century, abetted by the availability of human refuse in uncontrolled garbage dumps. This new food source greatly increased the survival of juvenile gulls. The large gulls displaced smaller gulls and terns from their traditional nesting sites. They also preyed on the eggs and chicks of these species. New landfill practices and alternative refuse management have reduced this food source in many urban areas; the populations of some gulls have begun to decline. Conservation measures include protecting colonies from direct exploitation (hunting of adults, egging), creation of suitable nesting space, stabilization of ephemeral nesting habitats, construction of artificial nesting islands or platforms, removal of predators (feral cats, large gulls, foxes), protection from other predators, and reduction of human disturbance at colony sites through sign-posting, fencing, or even
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wardening. More difficult, but equally important, is the protection of foraging sites and the prey base, which may involve fisheries management.
See also Beaches, Physical Processes Affecting. Network Analysis of Food Webs. Oil Pollution. Sea Level Change. Seabird Conservation. Seabird Foraging Ecology. Seabird Migration. Seabird Reproductive Ecology. Seabirds and Fisheries Interactions. Seabirds as Indicators of Ocean Pollution.
Further Reading Burger J and Gochfeld M (1996) Laridae (gulls). In: delHoyo J, Elliot A, and Sargatal J (eds.) Handbook of the Birds of the World, pp. 572--623. Barcelona, Spain: Lynx Ed. Gochfeld M and Burger J (1996) Sternidae (terns). In: del Hoyo J, Elliott A, and Sargatal J (eds.) Handbook of the Birds of the World, pp. 624--667. Barcelona, Spain: Lynx Ed. Grant PJ (1997) Gulls: A Guide to Identification. Vermillion, South Dakota: Buteo Books. Olsen KM and Larsson H (1995) Terns of Europe and North America. Princeton, NJ: Princeton University Press. Tinbergen N (1960) The Herring Gulls World. New York: Basic Books. Zusi RL (1996) Rynchopidae (skimmers). In: del Hoyo J, Elliott A, and Sargatal J (eds.) Handbook of the Birds of the World, pp. 668--677. Barcelona, Spain: Lynx Ed.
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LAW OF THE SEA P. Hoagland, J. Jacoby and M. E. Schumacher, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1481–1492, & 2001, Elsevier Ltd.
Introduction The law of the sea is a body of public international law governing the geographic jurisdictions of coastal States and the rights and duties among States in the use and conservation of the ocean environment and its natural resources. The law of the sea is commonly associated with an international treaty, the Convention on the Law of the Sea (UNCLOS), negotiated under the auspices of the United Nations, which was signed in 1982 by 117 States and entered into force in 1994. At present 133 States have signed and ratified UNCLOS; Canada, Israel, Turkey, USA, and Venezuela are the most prominent among those that have not ratified. This treaty both codified customary international law and established new law and institutions for the ocean. UNCLOS is best understood as a framework providing a basic foundation for the international law of the oceans intended to be extended and elaborated upon through more specific international agreements and the evolving customs of States. These extensions have begun to emerge already, making the law of the sea at once broader, more complex, and more detailed than UNCLOS per se. The law of the sea can be distinguished from two closely related bodies of law: maritime and admiralty. Maritime law is the private law relating to ships and the commercial business of shipping. Admiralty law, often used synonymously with maritime law, applies to the private law of navigation and shipping, in inland waters as well as on the ocean. The latter may also refer more parochially to the legal jurisdiction of specialized Admiralty courts. There may be important overlaps between the public international law of the sea and private maritime law, as may occur through the application of rules for vessel passage through a jurisdiction or the enforcement of domestic law in the ocean. The historical development of the law of the sea is sometimes traced back to a Papal Bull of 1493, which divided the world’s oceans between Portugal and Spain, thereby solidifying Spain’s claim to Columbus’ discovery of the New World. In the early
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seventeenth century, an important ‘debate’ took place between the Dutch jurist Hugo Grotius, who, in 1608, argued on the basis of natural law for freedom of the seas, and the English academic, John Selden, who argued in 1635 for the establishment of sovereign rights over areas of the ocean. In modern times, both regimes persist, although scientific and technological advances have combined to reduce that portion of the seas that is not subject to the authority of coastal States, and international rules have been developed to regulate many types of activities that occur beyond the reach of national jurisdictions. This article outlines the public international law of the sea, focusing mainly on UNCLOS. Important extensions of the UNCLOS framework are highlighted. The development of the law of the sea can be conceptualized as a tree with UNCLOS as its trunk. Its roots are historical customs, some centuries old, and agreements that emerged mostly after World War II. Its branches are customs, agreements, and soft law that is only now beginning to take shape. Six topical areas are covered: underlying principles, jurisdictions, fishery resources, mineral resources, marine science and technology, environmental protection, and dispute settlement.
Underlying Principles UNCLOS and its related agreements articulate certain distinctive, but closely related, principles of international environmental law. One of these, concerning sovereignty over resources, can be considered a general principle of customary international law. Others, including precautionary action, the common heritage of mankind, the duty to conserve the environment, sustainable development, and international cooperation, are just now emerging. These latter are philosophical concepts helping to shape the law of the sea that may one day achieve the status of general principles. Sovereignty over Resources
One of the most widely accepted norms of international environmental law is found in Principle 21 of the Stockholm Declaration of 1972. Its objective is to strike a balance between a State’s sovereignty and its responsibility to ensure that its activities and the activities of its citizens do not cause environmental harm to other States or to areas beyond national
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LAW OF THE SEA
jurisdiction. The UNCLOS rendering of Principle 21 reads: States have the sovereign right to exploit their natural resources pursuant to their environmental policies and in accordance with their duty to protect and preserve the marine environment.
Further, States shall take measures necessary to ensure that activities under their jurisdiction or control are so conducted as not to cause damage by pollution to other States and their environment, and that pollution arising from incidents or activities under their jurisdiction or control does not spread beyond the areas where they exercise sovereign rightsy
This principle applies to the actions of the citizens of a State within its territorial sea and exclusive economic zone, as well as on ships flying its flag, wherever they may steam. Precautionary Action
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threat to the environment. The principle is stated in connection with the Area (see next section on Jurisdictions), which, along with its resources, is defined explicitly in UNCLOS as the common heritage of mankind. Some commentators have argued, however, that the exploitation of the resources of the Area is still a high seas freedom, not subject to the common heritage principle. This latter interpretation may be particularly relevant for States that have not ratified UNCLOS. Environmental Conservation
UNCLOS specifies that: All States have the duty to take, or to cooperate with other States in taking, such measures for their respective nationals as may be necessary for the conservation of the living resources of the high seas.
Thus, a State whose nationals fish on the high seas is obliged to adopt conservation laws for its own citizens. The principle of obligatory environmental conservation under UNCLOS has influenced subsequent environmental agreements, including the 1985 ASEAN agreement through which its parties contracted to take measures to safeguard ecological processes, and soft law, including Chapter 17 (concerning the Oceans) of Agenda 21.
First applied to the marine environment in 1987 after the development of UNCLOS, the principle of precautionary action refines and strengthens Principle 21. During the last decade, it was incorporated increasingly into international agreements and soft law, such as the 1992 Rio Declaration and its accompanying Report on the United Nations Conference on Environment and Development (popularly known as Agenda 21). The articulation of the principle has been inconsistent, leading to varying interpretations in different contexts. In the context of marine pollution, a fair, but general, reading of the principle is that the release of substances thought to be potentially harmful should be regulated (or prohibited) prior to the establishment, according to scientific methods, of a causal link between the release and environmental damage. The principle implies a shift in the burden of proof from the pollutee or regulator, who previously had to prove that the release of a substance was harmful, to the polluter, who now must prove that it is not. The principle has an analogous interpretation in the fisheries context.
Sustainable development is another principle that emerged after the development of UNCLOS. It was articulated most clearly in the Rio Declaration, and it appears (referred to as sustainable use) in the 1992 Convention on Biological Diversity. As a general guiding principle, it implies economic or resource development in a way and at a rate such that the needs of both present and future generations can be met. Early conceptions of this principle appeared in UNCLOS, particularly with respect to the sustainable yield in fisheries, and it can be seen as closely related to the principles of environmental conservation and precautionary action.
Common Heritage
International Cooperation
Five principal elements characterize the common heritage of mankind: (1) common space areas are owned by no one but are managed by everyone; (2) universal popular interests have priority over national interests; (3) the economic benefits of natural resources exploited from the commons must be shared among all States; (4) the use of the commons must be limited to peaceful purposes; and (5) scientific research is permissible as long as there is no
In addition to the general obligation of members of the United Nations to cooperate in good faith with the organization and among themselves, UNCLOS expresses a particular need to cooperate to conserve the seas. The convention calls for international cooperation in the conservation and management of living and nonliving resources, the use of scientific study for the benefit of mankind, the peaceful settlement of all sea-related disputes, regulation of
Sustainable Development
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pollution, technology transfer to developing nations, and enforcement of all the provisions of the Convention. International cooperation is facilitated by international organizations, thus the Convention provides mechanisms to aid dialogue among member States. Some examples include: the International Sea Bed Authority, the International Tribunal on the Law of the Sea, and the Commission on the Limits of the Continental Shelf.
Jurisdictions The world’s oceans are divided into six basic zones in which the types and degrees of State jurisdiction vary. These zones are: the territorial sea, the contiguous zone, the exclusive economic zone (EEZ), the continental shelf, the high seas, and the Area. The seaward limits of the territorial sea, contiguous zone, and EEZ are defined in terms of distance from a baseline, which is essentially the waterline at low tide. The construction of baselines may follow any of several methods; in theory, the baseline might shift with changes in coastal geomorphology. The drawing of straight baselines is permitted across deeply indented coastlines or to connect islands along the coast of a State. (Islands, differentiated from mere rocks, must be capable of sustaining human habitation or an economic life of their own.) Baselines may not extend more than 24 nautical miles across the mouth of a bay. Territorial Sea
The territorial sea extends to a limit of 12 nautical miles from the baseline of a coastal State. Within this zone, the coastal State exercises full sovereignty over the air space above the sea and over the seabed and subsoil. A coastal State may legislate on matters concerning the safety of navigation, the preservation of the environment, and the prevention, reduction, and control of pollution without any obligation to make these rules compliant with international standards. Resource use within the territorial sea is strictly reserved to the coastal State. All States have the right of innocent passage through the territorial sea of another state, although there is no right of innocent air space passage. Innocent passage is considered moving through the territorial sea in a way that is not prejudicial to the security of the coastal State, including any stopping and anchoring necessary to ordinary navigation. Innocent passage implies two important limits to the power of coastal State jurisdiction in the territorial sea: (1) the obligation not to hamper, deny, or impair the right of innocent passage; and (2) the recognition
of innocent passage even in the case of vessel-source pollution as long as the pollution is not willful and serious. With notice, innocent passage may be suspended in specified areas of the territorial sea for security reasons. Even warships are to be accorded innocent passage (submarines must remain on the surface); however, in practice, many States require prior authorization for warships entering their territorial sea, and the law is unsettled here. Following the decision of the International Court of Justice in an infamous case in which Albania failed to notify Great Britain of the presence of underwater mines in the Corfu Channel, the coastal State must notify other States of its knowledge of navigational hazards. Regimes exist also for transit passage through international straits and archipelagic sea lanes passage in designated sea lanes through archipelagos, such as the Philippines. Contiguous Zone
The contiguous zone is a region adjacent to the territorial sea in which the coastal State may exercise control to prevent and punish infringement of its customs, fiscal, immigration, or sanitary laws. It may not exceed a distance of 24 nautical miles from the baseline. The coastal State may take action only with respect to offenses committed within its territory or territorial sea – not to those occurring within the contiguous zone or beyond. Although not sanctioned by UNCLOS, States such as India, Pakistan, and Yemen have asserted security jurisdiction in their contiguous zones. Such practices are becoming more widely accepted as customary international law. Exclusive Economic Zone (EEZ)
The EEZ is an area beyond and adjacent to a coastal State’s territorial sea to a limit of 200 nautical miles from the baseline. Within this zone, the coastal State may exercise sovereign rights over exploration, exploitation, conservation, and management of natural resources and other economic activities, such as the production of wind or tidal power. All States, whether coastal or land-locked, enjoy the right of navigation and overflight and the laying of submarine cables and pipelines within any EEZ. The coastal State alone, however, has the right to construct and operate artificial islands and other structural installations with accompanying 500 meter safety zones. Within the EEZ, the coastal State is primarily responsible for the conservation of living resources. The coastal State has the right to regulate both marine scientific research and pollution in the EEZ. It also has legislative and enforcement competence
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within its EEZ to deal with the dumping of waste from vessels and pollution from seabed activities. The practice of claiming an EEZ is one example of how UNCLOS has given rise to customary international law. The United States, for example, is not a party to UNCLOS but claims an EEZ that extends up to 200 nautical miles from its baseline. Canada has even adapted UNCLOS provisions to meet its needs for an exclusive fishing zone.
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immunity from the jurisdiction of any State other than their flag State. High seas fishing States have a duty to take conservation measures for their own nationals either alone or in cooperation with other nations. In instances of piracy, unauthorized broadcasting, slave trading, illicit drug trafficking, or statelessness, nonflag States may exercise enforcement jurisdiction. The Area
Continental Shelf
The continental shelf is geologically defined as the submerged prolongation of the land mass of the coastal State, consisting of the seabed and subsoil of the shelf, slope, and rise. It does not include the deep ocean floor. The significance of the continental shelf is that it may contain valuable minerals and shellfish. UNCLOS addresses the issue of jurisdiction over these resources by allocating sovereign rights to the coastal State for exploration and exploitation. The shelf has been defined as extending either to the edge of the continental margin or to 200 nautical miles from the baseline, whichever is further. Unlike the case of an EEZ, coastal States do not have to proclaim a continental shelf, but they must define its limits. Where the physical limits of the continental shelf extend beyond 200 nautical miles, the coastal State must delineate it, according to one of several formulas, using straight lines that do not exceed 60 nautical miles in length. A Commission on the Limits of the Continental Shelf makes recommendations to coastal States on matters related to the establishment of outer limits of the continental shelf where they extend beyond 200 nautical miles. High Seas
UNCLOS defines the high seas to be: All parts of the sea that are not included in the EEZ, the territorial sea, the internal waters of a State, or in the archipelagic waters of an archipelagic State.
On the high seas, all States enjoy freedoms of navigation, overflight, fishing, scientific research, the laying of submarine cables and pipelines, and the construction of artificial islands and installations. Because the high seas are open to all States, no State may attempt to subject any part of them to its sovereignty. Jurisdiction over ships on the high seas is reserved for the flag State. There must be a genuine link between the State and the ship that flies its flag, and States must fix their own conditions for granting nationality to ships and for registration. Warships and government vessels are accorded complete
The Area is defined as ‘the sea-bed and ocean floor and subsoil thereof, beyond the limits of national jurisdiction’. The Area has significance because of the occurrence of mineral resources, such as polymetallic nodules. Like the rest of the high seas, the Area and its resources are considered to be the common heritage of mankind. Each State must ensure that the activities of its own nationals are controlled, with the understanding that damage to the Area may entail State liability. An International Seabed Authority, established in Jamaica, regulates all activities in the Area, from marine scientific research to resource exploration and development. The Authority also has the right to conduct scientific research and to enter into research contracts. Finally, it enjoys the right to make rules and regulations preventing pollution to the marine environment and protecting natural resources. All installations are subject to these rules and regulations. Boundary Determinations
Several territorial and continental shelf boundaries were decided prior to the signing of UNCLOS, but there are many international boundaries that still must be drawn. In 1984, in the Gulf of Maine Case, the International Court of Justice decided the first combined EEZ and continental shelf boundary, between the United States and Canada. There appear to be no hard and fast rules for boundary determinations. Rationales for claims have ranged from historic uses to economic significance, leading the Court to decide most cases on the basis of equitable principles and relevant circumstances.
Fishery Resources The last half-century bore witness to significant growth in worldwide yields of marine fish stocks, starting at around 20 million metric tons in 1950 and peaking at 93 million metric tons in 1997. Although each fishery has its own unique characteristics, fisheries scientists now believe that, at the global level, aggregate yields have approached a natural limit. There are well-known examples of fisheries that have
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been exploited at inefficiently high rates, leading in some cases to severe stock depletion (e.g., north–west Atlantic cod). Evidence continues to mount of a shift from the exploitation of species at high trophic levels to those at lower levels, revealing a natural constraint to further expansion of wild harvests. Any increases in the production of seafood from the ocean and its value are likely to require both the implementation of more effective management measures that seek to optimize economic yields and the continuing development of husbandry (aquaculture). In the face of production trends and constraints in wild harvest fisheries, there is a critical need for the implementation of management measures that lead to sustainable yields. Although this need has been recognized for decades, it has rarely been achieved because of the difficulties of allocating shares of harvests across different groups in the face of limits to understanding the dynamics of intertwined ecological and environmental systems. As a practical matter, the international law of the sea relating to fisheries conservation provides only a crude framework within which to work. Domestic and regional institutions implement specific management measures within the broader context of this framework. Regional Fishery Management
Regional institutions were established as early as a century ago primarily for the purposes of conducting scientific research on fisheries and ecosystems that would lead, it was hoped, to recommendations for management (namely, the International Council for Exploration of the Seas in 1902). In the period since the end of World War II, these regional institutions proliferated. Today, more than 30 regional fishery bodies exist worldwide, most of which now strive to couple fisheries science to the active management of stocks that straddle the fisheries jurisdictions of multiple States or stocks that are located in part beyond any national jurisdiction (the so-called highseas and highly migratory stocks). Where stocks are actively managed, national quotas tend to be the instrument of choice, although enforcement problems are rife. Even with this institutional presence, at any time, dozens of fisheries conflicts are occurring between the nationals of different States. In the extreme, fishery conflicts have been known to escalate to the level of military intervention. The impetus for extending national territories that led to the basic jurisdictional zones codified in UNCLOS was driven by the perceived value of marine resources adjacent to coastal States. For fishery resources in the developed world, this value increased as demand expanded and technological
innovations reduced costs. In 1958, an international Convention on Fishing and Conservation of the Living Resources of the High Seas was signed, providing the basic framework that remains little changed to this day: local management coupled with the encouragement to cooperate internationally where nationals from different States prosecute the same fishery. According to the 1958 Convention, States were permitted to implement conservation and management measures for their own nationals fishing ‘high seas’ stocks adjacent to their coasts and were urged to cooperate with other States fishing there. Fishery Conservation Zones
One shortcoming of the 1958 Convention was that the geographic boundary defining the high seas was left undefined. This problem was rectified by UNCLOS, which permitted States to claim an EEZ within which they could exercise ‘sovereign rights’ over the exploitation of their natural resources, including fisheries. Several conditions were placed on this exercise of sovereign rights, but, in practice, these conditions are not seen as limiting. For example, States may determine the total allowable catch and are to manage EEZ fisheries at levels that can produce a maximum sustainable yield. However, management for maximum sustainable yield may be qualified at the State’s discretion by economic, environmental, ecological, or distributional reasons. These qualifications could be used as arguments for setting allowable catch, and thereby fishing effort, at levels either above or below those that might maximize sustainable yield. If a coastal State does not have the capacity, as measured by itself, to harvest its allowable catch, then, by agreement, it shall give other States, including landlocked and geographically disadvantaged States, access to any surplus. (This provision does not apply to sedentary shellfish stocks anywhere on the continental shelf.) In practice, the discretion accorded a State in determining fishing capacity and allowable catch implies that any surplus could be defined away easily. However, some States, notably Pacific Island States, have used these provisions to rent out their EEZ fisheries to the fleets of major distant water fishing nations, such as Japan. Coastal States within whose internal waters and EEZs anadromous fish (e.g., salmons) originate or catadromous fish (e.g., eels) spend the greater part of their life cycle are responsible for management of these species. Unless otherwise agreed to on a regional or an international basis, such species are to be fished inside the EEZ. Highly migratory species (e.g., tunas, billfishes, sharks, cetaceans) are to be
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managed through regional or international organizations to ensure conservation and to promote optimum utilization. Importantly, marine mammal conservation may be regulated more strictly within a coastal State’s EEZ than provided for by international rules. Straddling and High Seas Stocks
UNCLOS also provides a framework for straddling and high seas stocks. This framework has been elaborated further in a 1995 international Agreement on Straddling Fish Stocks and Highly Migratory Fish Stocks. Problems remain, however, including the specification of multiple and potentially mutually exclusive management objectives (e.g., maximize yield and minimize by-catch). Again, regional bodies are asked to undertake the tough job of operational management. States are encouraged to join existing or to establish new regional management institutions. However, where such bodies already exist, the basis for incorporating new entrants into decision making and for allocating to them a limited quota remain unclear.
Mineral Resources Ocean mineral resources, particularly offshore oil and natural gas, contribute significantly to worldwide supply. Offshore deposits now provide almost 10% of oil and 20% of natural gas production worldwide. Hard mineral deposits are much less important, although in some areas their production is meaningful to local economies. Tin has been produced for decades by dredging high-grade deposits located in the nearshore waters of Thailand and Indonesia. Diamonds are now profitably recovered off the coast of Namibia. Sulfur and salt are mined in conjunction with offshore oil production. Sand and gravel and calcium carbonate for use as a construction aggregate and to forestall beach erosion are dredged in many parts of the world. Other minerals on the continental shelves include phosphorite deposits and heavy mineral sands. Interest in the exploration of these latter occurrences continues, but these resources cannot yet be classified as economic reserves. Certain types of deep ocean mineral deposits are plentiful, including polymetallic nodules, ferromanganese crusts, and polymetallic sulfides. Much political effort was expended to establish in UNCLOS an international legal regime governing the exploitation of these classes of minerals. Although deep ocean resources are thought to be vast, the cost of recovery and processing, including the
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major risks of operating on the high seas, cannot now or in the foreseeable future justify their commercial exploitation. Continental Shelf Minerals
Because of the costs of operating in the offshore environment, much of the production of ocean minerals takes place in shallow, near-shore waters. Deep-water facilities, which at present are operational only for oil and natural gas, such as those in the North Sea, require very large or high-grade deposits to generate viable scale economies. Where production takes place within the territorial sea, the legal regime is well developed, differing little from domestic rules onshore. Consequently, the most significant legal provisions in the international law of the sea relating to mineral resources concern the establishment of a regime for the continental shelf. Production from seabed pools of oil and natural gas began at the turn of the century off the coast of California and in the Gulf of Mexico. But it was not until after World War II that an international legal regime governing the disposition of the resources of the continental shelf began to take shape. In 1945, US President Harry Truman issued a Proclamation asserting US jurisdiction and control over the Continental Shelf seabed and the natural resources of the subsoil. No seaward limit to the shelf was specified, although it was suggested that the shelf could be considered to extend to a depth of 100 fathoms (D183 meters). The Truman Proclamation (and its companion proclamation concerning fishery resources) helped set off a series of jurisdictional claims of varying geographic and legal coverages in Latin America, the Middle East, and elsewhere. In 1958, a Convention on the Continental Shelf entered into force, defining the continental shelf as an area adjacent to a State’s coast – but beyond its territorial sea – to a depth of 200 m. The adjacent coastal State could exercise sovereign rights over the exploration and exploitation of the natural resources of its continental shelf. Importantly, this jurisdiction could be extended ‘to where the depth of the superjacent waters admits of the exploitation’ of the natural resources. In this sense, jurisdiction could be expected to ‘creep’ with technological advance and changes in market conditions. With respect to ocean mineral development, the activity surrounding the legal regime for the deep seabed arguably has drawn attention away from a more important part of UNCLOS: the royalty provisions concerning the development of the continental shelf. UNCLOS provides that nonliving resource production occurring on that portion of the
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continental shelf extending beyond 200 nautical miles is subject to financial payments or contributions in kind to the International Seabed Authority, which is to share them equitably among the parties to UNCLOS. Payments begin at 1% of the value or volume of production in the sixth year of production. The payment increases at 1% a year until it reaches 7% in the 12th year, where it remains fixed. Deep Seabed Minerals
As the Continental Shelf Convention was being finalized, economic geologists and mining engineers began to examine more closely the potential for exploiting the vast deposits of polymetallic nodules occurring on the deep seabed. Polymetallic nodules are composed of a number of metals, including iron, manganese, nickel, copper, and cobalt. Recent economic analyses focus on nodules mainly as a nickel ore, with cobalt, copper, and, in some scenarios, manganese to be produced as by-products. Early analyses, conducted in the late 1950s and early 1960s, suggested that the nodule resource was commercially exploitable, while noting that there was no legal mechanism for allocating rights to areas thought to be so far offshore as to be beyond national jurisdiction. In 1967, the Maltese Ambassador to the United Nations, Arvid Pardo, called for an international agreement to prevent the national appropriation of the deep seabed, to establish the seabed and its resources as a common heritage of mankind, and to employ any resource rents for the development of poor nations. Although these basic principles were eventually incorporated into UNCLOS, their acceptance by the international community, especially by the developed West and the Soviet bloc, was not immediate. By 1970, however, the administration of US President Richard Nixon proposed a common heritage mining regime for an International Seabed Area, located beyond the 200 m isobath, that laid the basis for the UNCLOS negotiations. When UNCLOS was ready for signature in 1982, the deep seabed regime had become so extraordinarily complex and restrictive as to be unpalatable to some of the western market-oriented States. The common heritage principle was to be the centerpiece of the postcolonial new international economic order, through which the development of the world’s poorer States would be boosted by mandatory technology transfer and the promise of financial payments flowing from mineral royalties. This conception was made to appear realistic in light of predictions of world resource limits, such as those made by groups like the Club of Rome, and short-
term upward trends in metal commodity prices. Those States concerned about the effects on their own mineral sectors from seabed mine production were appeased in part with the promise of production limits. The final treaty provided for a parallel system of mining. Each pioneer investor (either a State or an industrial consortium sponsored by a State) would stake a mining claim and offer an additional claim of equivalent expected value to the International Seabed Authority’s Enterprise. The Enterprise would mine the parallel claims using technology transferred to it by the industry. Although the parallel system was a US proposal, in 1982 the incoming administration of US President Ronald Reagan would have nothing to do with the deep seabed mining provisions. The United States, Germany, and Great Britain, all with industrial interests in deep seabed mining, refused to sign the Convention, arguing that the nonseabed provisions reflected customary international law. Other developed States with seabed mining interests, including Japan, France, Canada, The Netherlands, Australia, and the Soviet Union, signed the Convention but delayed ratification. In lieu of the UNCLOS regime, a reciprocating States regime was organized by the West, permitting claims to the deep seabed to be staked and recognized among the participants to that agreement. The combination of the alternative regime, a steep decline in commodity prices in the 1980s, and delayed ratifications resulted in an agreement in 1994 to modify the deep seabed mining regime. Among other provisions, the revised UNCLOS deep seabed mining regime eliminated production controls and mandatory technology transfers, reduced license fees, and put the claims of miners registered under the reciprocating States regime on an equal footing with pioneer investors registered under the Convention.
Marine Science and Technology UNCLOS was the first international agreement to establish a regime for the conduct of marine scientific research in the ocean. The regime recognizes the right of a coastal State to control access to ocean areas under its authority for the study of the physical characteristics of the ocean and its natural resources. Although the regime has been characterized by some in the scientific community as unnecessarily burdensome and too discretionary, and although problems in obtaining permission for scientific research commonly arise, the regime has proven to be workable. Seeking permission to conduct marine scientific research in the EEZ or on the continental shelf of
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another coastal State requires careful advance planning and, frequently, close cooperation with the scientific community in the coastal State. The Convention recognizes that any State or competent international organization has the legal right to conduct marine scientific research. This right is conditioned only on the rights and duties of other States. Marine scientific research must be conducted for peaceful purposes, using appropriate scientific methods, and in such a way so as not to interfere unjustifiably with other legitimate uses of the ocean.
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participate or to be represented in the research project and must be given access to all data and samples, assessments of data, and preliminary and final project results. Unless the coastal State acts to withhold consent within four months of the request or requires supplementary information, or unless outstanding obligations on the part of the requesting State exist, the consent of the coastal State is deemed to have been implied, and the research project may proceed without an affirmative grant of consent. Technology Transfer
Consent Regime
Within its territorial sea, each coastal State has the right to regulate, authorize, or conduct marine scientific research, as a specific exercise of its sovereignty there. The conduct of marine scientific research in the territorial sea of a coastal State requires its express consent. Within its EEZ and on its continental shelf, each coastal State has the right to regulate, authorize, or conduct marine scientific research, as a specific exercise of its jurisdiction there. The conduct of marine scientific research in the EEZ or on the continental shelf of a coastal State requires its consent. A coastal State may not exercise its discretion to withold consent for research on the continental shelf beyond 200 nautical miles unless such research is proposed in areas that have been specifically designated by the coastal State for exploration or exploitation. All States have the right to conduct marine scientific research in the water column beyond a coastal State’s EEZ and in the Area. States seeking consent to conduct marine scientific research must provide detailed information about a proposed research project at least six months in advance. Although it is free to do so, a coastal State is under no obligation to grant its consent for scientific research in its territorial sea. Conversely, under normal circumstances, a coastal State is to grant consent for EEZ or continental shelf research, and it must establish rules so that requests for research are not delayed or denied unreasonably. However, coastal States are given considerable discretion to withhold their consent for EEZ and Continental Shelf research. Notably, consent may be withheld if a scientific research project is of significance for resource exploration or exploitation, involves drilling or the use of explosives or harmful substances, involves the construction, operation, or use of artificial islands; if the request for consent contains inaccurate information about the nature and objectives of the project, or if the requesting State has outstanding obligations from a prior research project. If consent is granted, the coastal State has the right to
As an element of the new international economic order, language encouraging marine technology transfer was incorporated into UNCLOS to accelerate the social and economic development of the developing States. However, the technology transfer provisions are mainly hortatory, promoting international cooperation and suggesting options for program development. Importantly, there is no obligation to transfer technology other than on fair and reasonable terms and conditions, respecting the rights and duties of holders, suppliers, and recipients of marine technologies.
Environmental Protection UNCLOS was designed, in part, to serve as the unifying framework for international law on marine environmental protection, which it does primarily by clarifying the rights and duties of States in this regard. It provides general goals and a few recommendations for combating all forms of marine pollution and environmental degradation but no specific pollution-control standards or required actions. To the extent that specific standards and requirements exist, they are set forth in other multilateral agreements that address either a particular form or source of pollution or a particular area of ocean space. In addition to creating new legal instruments elaborating rules pursuant to the general goals of UNCLOS, States are called upon to cooperate in notifying other countries of imminent threats of pollution, eliminating the effects of such pollution and minimizing the damage, developing contingency response plans, undertaking research programs, exchanging data, establishing appropriate scientific criteria for pollution-control rules and standards, and implementing and further developing international law relating to responsibility and liability for damage assessment and compensation. Although UNCLOS requires States to take measures against pollution from any source, it places
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particular emphasis on certain categories of pollutant substances and sources. States must take measures designed to minimize to the fullest extent possible: (1) releases of toxic, harmful, or noxious substances, especially those that are persistent, from land-based sources, from or through the atmosphere, or by dumping; (2) pollution from vessels; (3) pollution from installations and devices used in exploration or exploitation of the natural resources of the seabed and subsoil; and (4) pollution from other installations and devices operating in the marine environment. Among the pollutant sources enumerated in UNCLOS, only vessel discharges and dumping by ships and aircraft are currently subject to detailed standards and regulations at the global level. Vessel Discharges
The main instrument addressing operational discharges by vessels is the 1973 International Convention for the Prevention of Pollution from Ships, as modified by the 1978 Protocol thereto. Known as MARPOL 73/78, this treaty system includes five annexes containing regulations for the prevention of pollution by oil, by noxious liquid substances in bulk, by harmful substances carried at sea in packaged forms or in freight containers, portable tanks, or road and rail wagons, by sewage from ships, and by garbage from ships. The annexes covering oil and noxious liquid substances in bulk are mandatory for all contracting parties, but the others are optional. Ocean Dumping
Pollution by dumping includes the deliberate disposal at sea of wastes or other matter from vessels, aircraft, platforms, or other artificial structures, as well as the deliberate disposal of the vessels, aircraft, or structures themselves. Dumping is regulated at the global level under the 1972 Convention on the Prevention of Marine Pollution by Dumping of Wastes and Other Matter, also known as the London Convention. The Convention prohibits the dumping of certain hazardous materials and limits the dumping of other wastes or matter by requiring prior permits, including special permits for some materials according to criteria relating to the nature of the material, the characteristics of the dumping site, and the method of disposal. An important category of wastes not covered by the London Convention are those derived from the exploration and exploitation of seabed mineral resources. Under UNCLOS, such wastes remain subject to regulation by individual States for activities conducted in areas under their jurisdiction, while wastes resulting from activities in the Area beyond national
jurisdiction are subject to regulation by the International Seabed Authority. UNCLOS also mitigates a more general shortcoming of the London Convention – the fact that it has only 78 Contracting Parties representing just 68% of world merchant-marine tonnage. The main benefits of UNCLOS in this regard are that it clarifies the rights of coastal States to prohibit dumping in waters under their jurisdiction and requires all of its Contracting Parties to enact domestic measures that are at least as stringent as the London Convention requirements. Movement of Hazardous Wastes
Another global agreement, of relevance to both vessel-source pollution and dumping, is the 1989 Convention on the Control of Transboundary Movements of Hazardous Wastes and Their Disposal (Basel Convention). Under the Basel Convention, transboundary movements of hazardous or other wastes can take place only upon prior written notification by the exporting State to the States of import and transit, and each shipment of waste must be accompanied by a detailed movement document. Land-based Marine Pollution
Land-based marine pollution (LBMP), although it accounts for an estimated 80% of all contaminants entering the sea, is regulated only at the national level throughout most of the world, with the exception of six regional seas where multilateral agreements are in force. Adoption of a global treaty on LBMP was the object of intensive diplomatic effort from the mid-1980s until 1995, when 109 States adopted instead the nonbinding Global Programme of Action for the Protection of the Marine Environment Against Land-Based Activities. Among the main factors discouraging adoption of a binding global convention have been the largely disappointing results of the regional agreements and the fact that the causes and effects of LBMP operate primarily at regional or smaller geographic scales. The main arguments in favor of a global convention have centered on the pervasiveness and seriousness of the problem in virtually all regions of the world and the inability of developing countries and regions to address it effectively in the absence of a legal mechanism that provides for the transfer of relevant technologies and other forms of assistance from the developed world. Airborne Marine Pollutants
No multilateral agreements are in force whose primary purpose is the regulation of airborne marine
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pollutants, but such pollutants are included within the general scope of several regional agreements that address a broad range of marine pollution sources. Of these, only the agreements covering the Baltic, North-East Atlantic, and Mediterranean include any specific regulatory measures. In addition, the 1979 Geneva Convention on Long Range Transboundary Air Pollution provides for detailed regulation of emissions of numerous airborne pollutants by participating Northern Hemisphere countries. Although it does not target marine pollution directly, the Geneva Convention presumably provides indirect benefits to the marine environment. Persistent Organic Pollutants
Potentially among the most significant international instruments for controlling marine pollutants that are both land-based and airborne is a draft global convention slated for adoption in late 2000. Commonly known as the POPs Treaty, the agreement will regulate the production, sale, and use of initially one dozen persistent organic pollutants (POPs), most of them pesticides, whose characteristics include the tendencies to bioaccumulate in the marine food chain and to undergo long-range oceanic and atmospheric transport. Habitat and Ecosystem Protection
UNCLOS calls for States’ pollution-control measures to include measures to protect habitats and ecosystems, but it does not make an explicit call for cooperation in this regard or for ecosystem-based management of marine resources. UNCLOS thus leaves large marine ecosystems, which typically straddle two or more jurisdictional zones, subject to potentially conflicting management approaches and enforcement standards. Protection of marine habitats is provided under two major international treaties – the 1975 Convention on Wetlands of International Importance Especially as Waterfowl Habitat (Ramsar Convention) and the 1992 Convention on Biological Diversity – and under several Regional Seas protocols and other regional agreements. Protection of marine ecosystems is far less well developed in international law, no doubt in large part because ecosystem science and management are themselves comparatively new and undeveloped fields. This circumstance may also account for what some legal scholars consider to be an incoherent approach to ecosystem protection in UNCLOS. The lack of clarity as to the locus of authority to enforce ecosystem protections is uncharacteristic of UNCLOS, which otherwise exhibits an overriding concern with jurisdictional clarity in the balance it
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strikes between the competing interests of international navigation and the environmental protection concerns of coastal States. In general, UNCLOS limits the authority of States to enforce national and international environmental regulations where such authority conflicts with other principles established under the various legal regimes relating to different categories of ocean space. For example, coastal State authority to enforce national laws is subordinated to the right of innocent passage in the territorial sea; and on the high seas, only the flag State of an offending vessel has authority to enforce international environmental regulations, in deference to the principle of freedom of navigation. Because of such provisions, in the view of some environmentalists, UNCLOS does not provide the basis for full and effective protection of the marine environment, even if its entire agenda of elaborating agreements is eventually completed.
Dispute Settlement Following the UN Charter, which requires that all States settle their international disputes by peaceful means and without endangering international security, UNCLOS provides a binding framework for the peaceful settlement of sea-related disputes. The Convention stipulates that if States cannot resolve their disagreements peacefully on their own, they are to submit them to one of the following international bodies of their choice: (1) the International Tribunal of the Law of the Sea; (2) an arbitral tribunal constituted in accordance with Annex VII of the Convention; (3) a tribunal set up in accordance with Annex VIII; or (4) the International Court of Justice. International Tribunal of the Law of the Sea
The Tribunal applies the provisions of UNCLOS and other rules of international law in deciding disputes. Its decisions are final and must be complied with by parties to the dispute. The decisions have binding force only among the parties and with respect to their particular dispute. The jurisdiction of the Tribunal comprises all disputes between parties to the Convention and the agreement relating to the implementation of the deep seabed mining provisions. The Tribunal is called upon to settle three types of claims: (1) claims that application of the International Seabed Authority’s rules and procedures are in conflict with obligations of the parties; (2) claims concerning excess jurisdiction or misuse of power; and (3) claims for damages to be paid for failure to comply with conventional or contractual obligations. The Tribunal
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also has jurisdiction over disputes concerning the Area through a special Seabed Disputes Chamber and can conduct judicial reviews of the International Seabed Authority. It cannot, however, substitute its own decision or measure for that of the Authority or annul any underlying rule, regulation, or procedure established by it. In 1999, a dispute between Saint Vincent and the Grenadines and Guinea was one of the first to be settled by the Tribunal. The M/V Saiga, a vessel flying the flag of Saint Vincent and the Grenadines, had been pursued and arrested by the Guinean Navy in international waters south of Guinea’s EEZ because illegal bunkering was alleged to have taken place within Guinea’s EEZ. The Tribunal was charged with making a judgment on whether Guinea could apply its customs laws in an area beyond its territorial sea. Although the ship’s master was eventually found guilty on several counts, the Tribunal found that Guinea’s application of customs laws in its EEZ was contrary to UNCLOS.
authority is recognized by UNCLOS. Some disputes regarding the law of the sea have already been brought before the ICJ. An important difference of arbitration under the ICJ is that once States accept the court’s jurisdiction, under the Statute of the International Court of Justice, acceptance of the final decision in any case cannot be withdrawn once proceedings are underway. Another difference is that parties cannot select the members of the court who will be hearing the case. Although arbitration before the ICJ depends on the willingness of both parties to agree to and to participate in the arbitration, the Court is powerful enough to exert influence over parties to a dispute who refuse arbitration. For example, in the 1974 Fisheries Jurisdiction case, although Iceland did not appear before the court, the fact that the case made it to the level of international arbitration put considerable pressure on Reykjavik to comply with applicable rules of international law.
Future Prospects Annex VII Arbitration
When parties to a dispute do not select a specific type of arbitration under Article 287, an arbitral tribunal under Annex VII is automatically formed. Arbitral tribunals formed under Annex VII are five-member tribunals. Each party appoints one member and the remaining three must be approved by both parties and must be nationals of third-party States. Decisions are made by a majority vote of its members. Annex VIII Arbitration
Arbitration under Annex VIII entails the establishment of four lists of experts from which arbitral tribunals may be constituted to hear special cases. Each party to the convention may nominate two experts in each of the fields. The lists are then established and maintained by four different international institutions: (1) the Food and Agriculture Organization for fisheries; (2) the UN Environment Programme for marine environmental protection; (3) the International Maritime Organization for navigation and ocean dumping; and (4) the Intergovernmental Oceanographic Commission for marine scientific research. Five-member tribunals are set up by these institutions to perform fact-finding and settle disputes.
The law of the sea will continue to evolve as the rising worldwide population places greater pressures on the natural resources and ecological systems of the coastal ocean. Most of these pressures, without question, will be situated in the territorial seas and exclusive economic zones. For example, the continuing and growing releases of macronutrients, such as nitrogen, from agricultural operations in all States may have far-reaching and cumulative impacts on coastal environments. To the extent that marine science can unveil the complex physical and ecological links among national marine jurisdictions, the relevance and import of the international law of the sea will grow. Further, a consequence of the scientific portrayal of coupled ocean–atmosphere systems will draw the law of the sea more tightly into the fold of international environmental law, integrating the broader field, and thereby rendering the law of the sea less distinguishable as a selfstanding body of law.
See also Fishery Management. International Organizations. Manganese Nodules. Mariculture, Economic and Social Impacts. Maritime Archaeology.
Further Reading International Court of Justice
The International Court of Justice (ICJ) is an independent forum for dispute settlement that was established under the UN Charter and whose
Center for Earth Science Information Network (2000) www.ciesin.org. Charney JI and Alexander LM (1991) International Maritime Boundaries. Boston: Martinus Nijhoff.
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LAW OF THE SEA
Churchill RR and Lowe AV (1988) The Law of the Sea. Manchester: Manchester University Press. D’Amato A and Engel K (1996) International Environmental Law Anthology. Cincinnati: Anderson. Food and Agriculture Organization (2000) www.fao.org. Group of Experts on the Scientific Aspects of Marine Pollution (1990) The State of the Marine Environment. UNEP Regional Seas Reports and Studies No. 115. Nairobi: United Nations Environment Programme. International Maritime Organization (2000) Summary of status of conventions as at 31 July 2000. www.imo.org/ imo/convent/summary.htm.
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Joyner CC (2000) The international ocean regime at the new millenium: a survey of the contemporary legal order. Ocean and Coastal Management 43: 163--203. Schoenbaum TJ (1994) Admiralty and Maritime Law, 2nd end, Vol. 1. St Paul: West Publishing. United Nations Law of the Sea (2000) www.un.org/Depts/ los.htm. United Nations Secretary General (1989) Law of the sea: protection and preservation of the marine environment. Report of the Secretary General A/44/461.
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LEEUWIN CURRENT G. Cresswell and C. M. Domingues, CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia & 2009 Elsevier Ltd. All rights reserved.
Introduction The Leeuwin Current (LC) is unusual among subtropical eastern boundary currents because it flows poleward instead of equatorward. Its annual cycle starts as far north as 171 S off northwestern Australia in March when sea level there increases due to the arrival of an annual sea level pulse from the western tropical Pacific via the Indonesian seas. The LC, marked at the surface by a shallow, warm, lowsalinity core, flows down the sea level slope into a seasonally weakening equatorward wind stress to
Cape Leeuwin at 341 S (Figure 1). There it turns eastward and flows toward the Great Australian Bight, which extends from about 1251 to 1351 E. As was pointed out recently by Ridgway and Condie, the arrival of the LC at the pivot point of Cape Leeuwin in April/May is fortuitous because it is then driven eastward by winter westerly wind forcing. By the time it reaches Tasmania (Figure 2), in a much diluted form, and with some name changes, it has traveled 5500 km from northwestern Australia. In addition to its tropical source, it receives contributions from the subtropical Indian Ocean to the west of Australia and from the subantarctic waters south of Australia. Off Perth, 321 S, its maximum poleward geostrophic transport is 5 Sv (5 106 m3 s 1) during June–July. It is strongest during La Nin˜a years. The LC is an important factor in the life cycles of many commercially important marine creatures
20
NW Cape 200 m Western Australia
25
Abrolhos Islands Capes Current (summer)
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Great Australian Bight ch Ar erch ch e ip el ag o
Latitude
Shark Bay
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Leeuwin Current
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120 Longitude
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Figure 1 A schematic picture of the surface currents off southwestern Australia. The LC is marked by the broad gray arrow and flows year-round, being strongest in winter. The Capes Current (long black arrow on the continental shelf) is driven by southerly windforcing in summer. Subtropical water (open arrows) is entrained into the LC system from the west, as is subantarctic water (black arrows) from the south. Two eddies that have separated from the main current are shown.
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Figure 2 A composite satellite sea surface image from several days in Jun. 1994. There is a 51 51 latitude–longitude grid; land is gray; the shelf edge is marked by a thin black line; clouds are white; and the temperature scale (1C) appears at the top of the image. The warm LC (false color red) flows southward to the SW corner of Australia (Cape Leeuwin) and then eastward. There are eddies south of the LC. Copyright 2001, CSIRO Marine Research, Hobart.
along the Western Australian seaboard. Annual fluctuations in the numbers of western rock lobsters (Australia’s most important single species fishery) recruited to the coastal reefs can be linked to the strength of the LC, which in turn can be related to El Nin˜o/Southern Oscillation (ENSO). Arripus trutta, Australian salmon, migrate westward across the Great Australian Bight to spawn north of Cape Leeuwin in autumn; the seasonally strengthening LC then carries their eggs and young toward the Great Australian Bight. Young southern bluefin tuna Thunnus maccoyii that hatch between NW Australia and Indonesia commence their Southern Hemisphere odyssey with partial assistance from the LC; they are found south of WA as 1-year-olds. Corals are found farther south than otherwise might be expected; the southernmost hatchery for the red-tailed tropic bird Phaeton rubricauda is found near Cape Leeuwin; and other tropical marine flora and fauna are carried to the Great Australian Bight.
History The evidence for an LC flowing from NW Australia south to Cape Leeuwin and then across toward Tasmania came piece by piece over almost two centuries. In May 1803, Matthew Flinders’ vessel, Investigator, was set to the east between Cape Leeuwin
(1151 E) and several degrees of longitude eastward at a little over 0.5 m s 1. From there to the Recherche Archipelago (1231 E) he found the current to increase with distance from the coast. Late in the nineteenth century, Saville-Kent suggested that tropical waters reached the Abrolhos Islands (B291 S) because of the warm waters and tropical marine flora and fauna, including corals, which he observed. Early in the twentieth century, Dakin learnt of a southward winter current inshore of the Abrolhos Islands from fishermen who also reported that it strengthened when northerly winds blew. In 1921, Halligan published a chart of the currents around Australia, showing how a cold and heavy Southern Ocean current approached the southwest coast of Australia and dipped beneath a warm southerly drift down the Western Australian coast. From Cape Leeuwin, he reported ‘‘a warm southerly and easterly surface current’’ with speeds of 0.2 m s 1. This fact must have come from ships that were close enough to the coast to measure their drifts with any precision – and therefore inshore of the strongest part of the LC. Schott, in his book in 1935, showed charts of temperatures and currents for the world, with warm water rounding Cape Leeuwin and flowing eastward, more or less as we know the LC today. A decade later, Rochord started the continental shelf sampling station at the 60-m isobath off Rottnest Island (321 S); the station continues to be occupied every few weeks.
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The time series of salinity, when combined with inferences from drift bottle releases and recoveries, suggested that low-salinity tropical water flowed southward to arrive at the station in autumn and winter and that there was a northward flow of highsalinity water in summer. In hindsight we recognize these as the Leeuwin and Capes Currents, respectively – in winter the LC spreads onto the continental shelf, while in summer it is seaward of the Rottnest station. Phytoplankton also provided clues: among them, Wood in 1954 postulated a current from Cape Leeuwin to Tasmania based on the continuity of findings of certain species of dinoflagellates. In 1969, Hamon reported that there was an eastward flow toward the shelf north and south of the Abrolhos Islands. Gentilli in 1972 was the first to see the importance of the throughflow of tropical Pacific Ocean
9 Apr.
waters into the NE Indian Ocean: he reported that these waters flowed southward against WA down to the latitude of Rottnest Island (321 S). Andrews in 1977 confirmed Hamon’s eastward inflow and found that it turned southward near the continent, spawned eddies, and turned eastward at Cape Leeuwin (cf. Figure 1). We now recognize this as the salty, subtropical input to the LC. In the same decade, Kitani reported a southward-flowing warm, low-salinity current between the Abrolhos Island and Fremantle as well as cyclonic and anticyclonic eddies. When satellite-tracked drifters were first released off WA in the 1970s (Figure 3), they revealed that the LC not only flowed southward along the continental slope to the Cape Leeuwin vicinity, but that it also rounded the cape, on one occasion quickly accelerating to 1.5 m s 1, and then flowed eastward at
Abrolhos Islands Geraldton
4 Apr.
Six Buoy tracks 23 March − 3 June, 1976
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Continental shelf
30° S
Buoy A Buoy B Buoy C Buoy D Buoy E Buoy F
23 Mar.
7 Apr.
9 May 29 Apr.
Perth
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Esperance Cape Naturaliste Cape Leeuwin
8 Apr.
3 Jun.
19 May
3 Jun.
22 May
Albany
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35°
30 May
10 May 16 May
14 May 29 Apr. 11 Apr.
30 Mar. 115° E
11 May 20 May
7 Apr. 28 Mar. 120°
Figure 3 The paths of satellite-tracked drifters during autumn when the LC increases to its maximum strength.
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LEEUWIN CURRENT
speeds up to 1.8 m s 1 toward the western side of the Great Australian Bight. The drifters showed that the LC meandered and formed eddies; for reasons that are not clear, the drifters most frequently went into cyclonic eddies, but there was one example of two of them moving from near the shelf edge off Perth out into an anticyclonic eddy (no. 7 in the figure) that appeared to be in the process of forming. The current south of WA, as indicated by the drifters, had small (B30 km) southward (anticyclonic) meanders separated by B200 km. At the locations of the southward meanders there could be transient offshoots out to sea that took the drifters into cyclonic eddies, where they remained trapped for some weeks. In some instances, the drifter motions showed the cyclonic eddies to drift westward at about 7 km d 1. The drifters decelerated when they entered the cooler cyclonic eddies and accelerated when they reentered the warmer LC. The groundings of several drifters along the south coast suggested that the LC could spread across the continental shelf. Some drifters came from the south to suggest that subantarctic waters were carried northward to interact with the LC. Off the western coast of WA the drifters also showed northward wind-driven flow on the continental shelf in summer, when the northward wind stress is strongest. This northward flow has been called the Capes Current and it includes waters upwelled off SW Australia (Figure 4). Because of the significance of Cape Leeuwin as the pivot point for the change from southward flow down WA to eastward flow to the GAB, the LC was given its name by Cresswell and Golding in 1980. The Leeuwin was a VOC (Dutch East India Company) vessel that explored the coastline in 1622 en route to Batavia. The first satellite sea surface temperature (SST) images in the late 1979s of the LC south of WA showed a number of small offshoots, with one being linked to an anticyclonic eddy so that the LC waters were carried over 200 km southward. The eddies and offshoots are now recognized to be a common feature of the LC. The arrival of the tropical waters of the LC at Cape Leeuwin in the austral autumn is a dramatic event: satellite images in 1987 (Figure 4) showed that the leading edge of the LC progressed southward and then eastward at about 0.25 m s 1 between March and April. In step with the arrival of the LC, there were ‘decreases’ in the northward winds off the west coast and ‘increases’ in the sea level. The 6 March image showed cool near-shore waters westward from 1181 E to Cape Leeuwin from upwelling driven by southeasterly winds. These were carried northward around Cape Leeuwin to
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contribute to the wind-driven Capes Current on the continental shelf (the green plume that crosses 331 S). The image for 15 April showed the northward intrusion of cold subantarctic water near 1181 E. The intrusion turned to the east to run alongside the LC. The existence of such intrusions could, in retrospect, also be inferred from the tracks and temperatures of the 1970s drifters.
Water Masses off Western Australia The water masses can largely be identified by their salinities, shown in the schematic diagram (Figure 5) drawn from several ship surveys. The diagram excludes continental shelf waters and currents such as the Capes Current. We note that the various water masses are found in layers variously atop, below, and alongside each other; that they move independently; and that mixing occurs across their boundaries. In the northwest (top left) is the inflow of nearsurface, low-salinity, warm tropical water (TW; blue; o34.5). This tropical input to the LC becomes saltier and cooler through mixing as it progresses southward (note that salinities lower than 35.5, that is, yellow, are found near the continent almost as far south as Cape Leeuwin). Air–sea fluxes play a small part in the temperature and salinity changes observed along the poleward path of the mean jet of the LC. Most of the changes seem to be due to eddy fluxes. We have marked the LC as a light blue semitransparent arrow running southward to Cape Leeuwin and then eastward to the Recherche Archipelago and then beyond to the Great Australian Bight (top right). In reality, the LC flow is broken up by meanders and eddies. In the west is the Hamon–Andrews inflow of subtropical water (STW) and South Indian Central Water (SICW), ranging from 435.5 (red) near the surface (due to an excess of evaporation over precipitation) down to the low-salinity (blue; o34.5) Antarctic Intermediate Water (AAIW) at more than 600-m depth. The salty STW (red) is entrained to become an important part of the LC and it is carried around Cape Leeuwin both near the surface and near the continental shelf edge, so that the LC changes from being a low-salinity current relative to its surroundings west of Australia to being a high-salinity one south of Australia. Salty STW and the lowersalinity central water sink beneath the LC to become part of the northward-flowing Leeuwin Undercurrent (LUC) between c. 300 and 700 m. To the south of WA the green arrow marks the Subantarctic Mode Water (SAMW) that results from deep winter mixing at about 451 S. This can be traced as it moves northward off WA by its relatively high
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Figure 4 Satellite images 6 weeks apart showing the energetic nature of the arrival of the warm waters of the LC. There is a 11 11 latitude–longitude grid; land is black; the shelf edge is marked by a thin black line; clouds are white; and the temperature scale (1C) is at the bottom.
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LEEUWIN CURRENT
TW
NW
rn ste lia We stra Au
pe Ca
ay
ho
rol
Ab
la s Is
Leeuwin Current
e rch o he elag c Re chip Ar
s
nd
kB
ar Sh
0
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rth
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n
wi
eu
e .L
C
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1000 m LUC
STW
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36
Salinity Figure 5 A schematic diagram of the salinity structure down to 1000 m drawn from several ship surveys. The boxes are 1000-m deep and about 200 km wide. The view is from above the southern Indian Ocean and the map is to orient the reader.
oxygen values as compared with the deeper component of the SICW from the west that has the same temperature and salinity values. The blue arrow indicates the AAIW originating at the surface much farther south. The white arrows attempt to show the undercurrent in the south that flows westward to Cape Leeuwin beneath the eastward-flowing LC and then joins – or becomes – the northward-flowing LUC.
The Leeuwin Current and Its Eddies The LC path southward from NW Cape to Cape Leeuwin is a meandering one and it spawns and interacts with anticyclonic and cyclonic eddies. The anticyclonic eddies can be 200 km across and they are more akin to eddies in western boundary systems, such as the EAC, having edge speeds of as much as 1m s 1, long lifetimes (41 year), and mixing to several hundred meters depth in winter. The eddies can be followed with satellite altimetry measurements as they drift into the central Indian Ocean. Both cyclonic and anticyclonic eddies can be tracked, with the former having a slight poleward bias to their drifts and the latter an equatorward one.
In Figure 6, the LC is seen near the upper continental slope where the salinity and temperature structure dip downward. The current has a subsurface maximum, probably because of strong southerly winds at the time. The structure farther out to sea reveals the upward doming associated with a cyclonic eddy. The high-salinity water is subtropical surface water from farther westward, although on the shelf insolation and evaporation have warmed the waters and made them more saline. South of Western Australia the LC can be particularly strong (Figure 7) and the situation with eddies is different: weak anticyclonic eddies, that form as far east as Tasmania, drift westward to encounter the LC, and the continental slope near the Recherche Archipelago. There they steer the warmer LC waters offshore and around them (Figure 8) and over their tops, with the result that they strengthen significantly. The two eddies in the image drifted westward out into the Indian Ocean. The warm waters of the LC spread inshore across the shelf, although the current would be strongest just seaward of the shelf edge. Ship measurements across an offshoot joining the LC to an anticyclonic eddy revealed a trough-shaped structure with southward flow of 0.5m s 1 on the
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Marmion section, Dec. 1994 36
0
Depth
>35.9 –100
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Figure 6 Ship measurements of salinity (practical salinity units) and temperature (1C) with a CTD, and north–south current with an acoustic Doppler current profiler (ADCP) west of Perth (321 S) in Dec. 1994.
Sea surface temperature
20°
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18° Salinity
16° XBT:
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*
*
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*
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10 km Figure 7 A north–south transect across the LC south of Western Australia near 1171 E in June 1987. Upper panel: surface temperature and salinity across the Leeuwin Current. Lower panel: current speed from acoustic Doppler current profiler (ADCP) measurements, the temperature structure from expendable bathythermograph (XBT) casts, and Richardson numbers calculated from the current and temperature structure. Dark shading, o0.25 (the region of most active overturning); light shading, o0.5. The current speed reached 1.5 m s 1. Reproduced from Cresswell GR and Peterson JL (1993) The Leeuwin Current south of Western Australia. Australian Journal of Marine Freshwater Research 44: 285–303.
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LEEUWIN CURRENT
14
15
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NOAA 1 1 SST 07 Jun 1994 0832Z Figure 8 An SST image showing the sea south of WA between Cape Leeuwin (left) and the western Great Australian Bight (top right). There is a 1 1 1 1 latitude–longitude grid; land is black; the shelf edge is marked by a thin black line; clouds are white; and the temperature scale (1C) is at the top. Copyright 2000, CSIRO Division of Marine Research, Hobart.
western side and the reverse on the eastern side (Figure 9). Cyclonic eddies form on the western sides of the seaward offshoots. The eddies drift westward, with the anticyclonic ones repeatedly interacting with the LC until they are west of the longitude of Cape Leeuwin. At times there can be three anticyclonic eddies between the Recherche Archipelago and Cape Leeuwin. Individual eddies have been followed using satellite altimetry for over a year from their first interaction with the LC at the Recherche Archipelago to beyond Cape Leeuwin and out into the Indian Ocean. The structure of an anticyclonic eddy south of WA is shown in Figure 10. The north and south current speed maxima were 0.5 m s 1. The eddy vertical structure suggested that it had mixed to over 300-m depth during the winter and then, with the onset of summer, had been warmed near the surface. The depression of the water structure extended at least to 1000 m. The eddy was low in oxygen and nutrients as compared with the surrounding waters. This contrasts with the anticyclonic eddies formed by the LC west of
Australia: there the formation process entrains relatively nutrient-rich waters from the continental shelf.
The Leeuwin Current on the Continental Shelf An acoustic Doppler current profiler (ADCP) moored a few meters above the bottom at the 70-m isobath out from Perth gave hourly data at 4-m depth intervals for little over an year (Figure 11). In December, there was strong northward alongshore flow due to the winddriven Capes Current that lowered the temperature at the instrument by 21C. The onset of the LC was quite sudden in late March and it was present in the lower half of the water column until the end of the record. Nearer the surface the LC flow was often weaker, perhaps due to those waters being more susceptible to northward wind forcing from passing storms. The lower shelf waters, particularly from April to August, were drained out across the topography with the offshore flow peaking at over 0.1 m s 1 in May. Whether
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20°
Sea surface temperature
18°
Salinity
35.9 35.8 35.7 35.6
16° XBT:
48
*
49
Southward
*
50
51
Northward
*
*
52
*
>0.7
–0.5 0.5
50 Depth (m)
0.0 19° 100
18° 17°
150 16° 200
25 km Figure 9 An east–west transect south of Western Australia at 1171 E in Jun. 1987 across an offshoot connecting the LC to an anticyclonic eddy. Upper panel: surface temperature and salinity. Lower panel: the north–south current components from acoustic Doppler current profiler (ADCP) measurements, the temperature structure from expendable bathythermograph (XBT) casts, and Richardson numbers calculated from the current and temperature structure. Dark shading, o0.25; light shading, o0.5. Reproduced from Cresswell GR and Peterson JL (1993) The Leeuwin Current south of Western Australia. Australian Journal of Marine Freshwater Research 44: 285–303.
Offshoot eddy at 37° S, Dec. 1994, stns. 32:42
Depth
0
>35.6
35.5
–200 35 –400 Salinity 115
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117
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119
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16 14
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–0.4
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Longitude Figure 10 A transect of salinity (practical salinity units), temperature (1C), and north–south velocity component (m s 1) from 381 S, 1151 E to 371 S, 1201 E across an anticyclonic eddy south of WA in Dec. 1994. The velocity component plot has been limited to the 300-m depth reached by the acoustic instrument taking the measurements.
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(a)
Monthly PVDs Nov. 00 to Nov. 01 for selected depths 800
600 Nov.
Dec.
Jan.
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Progression along shore (km)
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Progression across shore (km) (b)
Alongshore NCEP wind stress at 32° S, 114° E
Wind stress (N m−2)
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Figure 11 (a) The monthly progressive vector diagrams, or inferred water trajectories, for selected depths (every fourth depth bin) from Nov. 2000 to Nov. 2001. Only data from every fourth depth bin are shown. The trajectories are arranged in columns, with successive 200-km offsets, and by row, for the different depths. The axes are across and along the bottom topography. (b, c) The alongshore wind stress near the mooring site and the temperature recorded by the near-bottom instrument moored at the 70-m isobath.
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this was due to an Ekman bottom boundary layer beneath the LC, or to near-shore waters cooling and becoming denser and moving out across the shelf, or both, is not clear. Note that the near-bottom temperaturve (Figure 11(c)) was highest in the austral autumn/winter (March/July) due to the influence of the warm LC waters from the north.
See also Current Systems in the Indian Ocean. Current Systems in the Southern Ocean. Drifters and Floats. East Australian Current. Ekman Transport and Pumping. El Nin˜o Southern Oscillation (ENSO). Energetics of Ocean Mixing. Inverse Modeling of Tracers and Nutrients. Meddies and Sub-Surface Eddies. Mesoscale Eddies. Moorings. Satellite Remote Sensing of Sea Surface Temperatures. Sea Surface Exchanges of Momentum, Heat, and Fresh Water Determined by Satellite Remote Sensing. Single Point Current Meters. Upper Ocean Mixing Processes. Upwelling Ecosystems. Water Types and Water Masses.
Further Reading Andrews JC (1977) Eddy structure and the West Australian Current. Deep-Sea Research 24: 1133--1148. Cresswell GR and Golding TJ (1980) Observations of a south-flowing current in the southeastern Indian Ocean. Deep-Sea Research 27A: 449--466. Cresswell GR and Griffin DA (2004) The Leeuwin Current, eddies and sub-Antarctic waters off south-western Australia. Marine and Freshwater Research 55: 267--276. Cresswell GR and Peterson JL (1993) The Leeuwin Current south of Western Australia. Australian Journal of Marine Freshwater Research 44: 285--303. Domingues CM, Wijffels SE, Maltrud ME, Church JA, and Tomczak M (2006) Role of eddies in cooling the Leeuwin current. Geophysical Research Letters 33: L05603 (doi:10.1029/2005GL025216). Feng M, Meyers G, Pearce A, and Wijffels S (2003) Annual and interannual variations of the Leeuwin Current at 321S. Journal of Geophysical Research 108(C11): 3355 (doi:10.1029/2002JC001763).
Fieux M, Molcard R, and Morrow R (2005) Water properties and transport of the Leeuwin Current and eddies off Western Australia. Deep-Sea Research I 52: 1617--1635. Godfrey JS and Ridgway KR (1985) The large-scale environment of the poleward-flowing Leeuwin Current, Western Australia: Longshore steric height patterns, wind stresses and geostrophic flow. Journal of Physical Oceanography 15: 481--495. Holloway PE and Nye HC (1985) Leeuwin Current and wind distributions on the southern part of the Australian North West Shelf between January 1982 and July 1983. Australian Journal of Marine and Freshwater Research 36: 123--137. Middleton JF and Platov G (2003) The mean summertime circulation along Australia’s southern shelves: A numerical study. Journal of Physical Oceanography 33: 2270--2287. Morrow R, Birol F, Griffin D, and Sudre J (2004) Divergent pathways of cyclonic and anti-cyclonic ocean eddies. Geophysical Research Letters 31: L24311 (doi:10.1029/ 2004GL020974). Morrow R, Fang FX, Fieux M, and Molcard R (2003) Anatomy of three warm-core Leeuwin Current eddies. Deep-Sea Research II 50: 2229--2243. Pearce A and Pattiaratchi C (1999) The Capes Current: A summer countercurrent flowing past Cape Leeuwin and Cape Naturaliste, Western Australia. Continental Shelf Research 19: 401--420. Pearce AF and Phillips BF (1988) ENSO events, the Leeuwin Current, and larval recruitment of the western rock lobster. Journal Du Conseil 45(1): 13--21. Ridgway KR and Condie SA (2004) The 5500-km long boundary flow off western and southern Australia. Journal of Geophysical Research 109: C04017 (doi:10.1029/2003JC001921). Schodlok MP, Tomczak M, and White N (1997) Deep sections through the South Australian Basin and across the Australian–Antarctic discordance. Geophysical Research Letters 22: 2785--2788. Smith RL, Huyer A, Godfrey JS, and Church JA (1991) The Leeuwin Current off Western Australia, 1986–1987. Journal of Physical Oceanography 21: 323--345. Waite AM, Thompson PA, Pesant S, et al. (2007) The Leeuwin Current and its eddies: An introductory overview. Deep-Sea Research II 54: 789--796.
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LONG-TERM TRACER CHANGES F. von Blanckenburg, Universita¨t Bern, Bern, Switzerland Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1492–1503, & 2001, Elsevier Ltd.
Introduction Ocean tracers that record long-term changes preserve certain water column information within the sediment. This information comprises (1) the tracers’ fluxes in the past, such as erosional input from the continents, hydrothermal activity at mid-ocean ridges, input of extraterrestrial material, or carbonate recycling; (2) the distribution of water masses in the past and the state of the past global thermohaline circulation. Inorganic isotope tracers whose isotope ratios are modified by radioactive decay in their source materials are ideally suited for these studies. Their original water column values can be measured in materials such as biogenic carbonates, ferromanganese crusts and nodules, and the authigenic phase of deep-sea sediments. Studies of tracer fluxes in the past are favored by those tracers whose residence time in the ocean (t, defined below) is long relative to the turnover time of the thermohaline circulation (1500 y), such as Sr and Os. Tracers of which t is of the order of, or shorter than the oceans turnover time (Nd, Hf, Pb, Be) offer the ability to label water masses isotopically. In this case, long-term isotope changes of these intermediate-t tracers are potentially caused by variations of the thermohaline circulation. However, secular variations of these isotope tracers can also be caused by regional variations in these tracers’ fluxes, mostly resulting from changes in weathering. It is not always straightforward to distinguish between these two causes of tracer variations. Certainly the globally uniform seawater isotope evolution of Sr, Os, and potentially also Be, offer excellent tools for isotope stratigraphy on long (My) timescales.
Definitions and Concepts Long-term tracers are those elements whose isotopic compositions provide information on the physical and chemical state of the oceans on timescales of several thousands of years to millions of years (My). For example, paleo-oceanographers aim to
reconstruct past water mass distributions and the mode of the thermohaline circulation. For this purpose it would be desirable to reconstruct past oceanographic water mass characteristics such as salinity, temperature, silica, or phosphorus content from the sedimentary record. Similarly, the reconstruction of the past land–sea transfer of certain tracers is desirable in order to reconstruct changes in the weathering history of the continents. However, these present-day tracers are usually not conserved in the sedimentary record. Even if they were precipitated chemically and stored in sediments, their changes in concentration as measured in a sedimentary column back through time cannot be directly related to past water mass properties. This is because the tracers’ concentrations in sediments depend on factors such as sedimentation rate, diagenesis, partitioning into a certain phase, uptake by organisms, and additions of the same element by detrital hemipelagic or aeolian material. Therefore ocean chemists make use of proxy tracers which are not routinely analyzed in surveys of present-day water masses, because their measurement presents a considerable effort compared with tracers such as salinity and temperature. However, their characteristics can be directly related to those well-known oceanographic seawater tracers. The conditions that need to be met for an element to be of use as a proxy tracer are that (1) it conserves a characteristic chemical or isotopic property when transferred from the water column into the sediment; (2) the elements or their isotopic composition can be extracted from the sediment; (3) the age of the sediment is known so that changes of the tracer over time can be reconstructed. One such proxy makes use of element ratios. For example, the ratio of Cd to Ca in foraminiferal tests is a proxy for the PO4 content of the past water mass in which the foraminifera formed and therefore provides information on the past thermohaline circulation. This tracer is explained in detail in the article on trace elements in foraminiferal tests. Similarly, the ratio of the intermediate uranium decay products 231Pa and 230Th, measured in bulk sediment, may under certain conditions provide information on the advection of water masses in the overlying water column in the past (see Cosmogenic Isotopes and Uranium-Thorium Series Isotopes in Ocean Profiles). Isotope ratios are ideally suited as long-term proxy tracers. Some of these isotope ratios are characteristic of certain seawater properties and
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LONG-TERM TRACER CHANGES
can be measured in sediments, regardless of the actual tracer partitioning, concentration, or location of precipitation. A property describing the behavior of a tracer in sea water is the residence time t. If the tracer’s fluxes in and out of an ocean basin are invariant with time, the tracer is at steady state and the residence time can be calculated from the tracer’s ocean inventory: t¼
Inventory Inventory ¼ Fluxin Fluxout
A tracer suitable as a water mass tracer has a short global residence time (t) relative to the ocean’s mixing time. This ensures that isotope ‘fingerprints’ characteristic of a certain water mass are prevented from being completely dispersed by the global thermohaline circulation. While the global ocean mixing time is difficult to assess, a meaningful quantity is the time it takes for one turnover of the global deep water circulation, which is c. 1500 y. Tracers with t of this order have the potential to preserve distinct water mass labels. It can be assumed that a conservative (i.e. nonreactive) tracer would be almost perfectly homogenized within 10 000–20 000 years. Tracers with t in excess of this period will only Table 1
record changes in the global flux of this tracer, regardless of the water mass, the location of the input, or the location of the samples taken.
Isotope Tracers Used Much use is made of the stable isotopes of carbon as a paleo-water mass isotope ‘fingerprint’. The 13C/12C ratio in the tests of foraminifera depends on the relative position of the overlying water mass within the thermohaline circulation system. However, these isotope ratios are modified during the incorporation into organisms, depend on availability of nutrients, and like the isotopes of oxygen, also depend on seawater temperature. (These tracers are dealt with in the relevant articles; please refer to the See also section.) Isotope ratios of inorganic trace metals which are the topic of this chapter are not modified when incorporated into the sediment (note that some minor isotope fractionation might occur on incorporation into the sediment, but usually such shifts are either smaller than analytical precision or they are removed by the internal correction procedures of the techniques used). The variation in isotope ratios only varies
Long-term isotope tracers currently in use
Tracer
Isotopes
Sources
Average deep-water concentration
Global deep-water residence time
Strontium (Sr)
87
Sr (stable) ’87Rb (T1/2 ¼ 48.8 Gy) 86 Sr (stable, primordial)
7.6 mg g1
2–4 My
Osmium (Os)
187
10 fg g1
8000–40 000 y
Neodymium (Nd)
Nd (stable) ’ 147Sm (T1/2 ¼ 106 Gy) 144 Nd (stable, primordial) 176 Hf (stable) ’ 177Lu (T1/2 ¼ 37.3 Gy) 177 Hf (stable, primordial)
Mostly chemical weathering of the continental crust and carbonates Hydrothermal solutions from midocean ridges Dissolution of marine carbonates Erosion of the continental crust (chemical weathering important) Leaching of abyssal peridotites Cosmic dust and spherules Erosion of the continental crust
4 pg g1
B1000–2000 y
Erosion of the continental crust
0.18 pg g1
B1000–2000 y?
Hydrothermal solutions at mid-ocean ridges Erosion of the continental crust
1 pg g1
40 y (Atlantic)
Hafnium (Hf)
Lead (Pb)
Be Beryllium
Os (stable) ’ 187Re (T1/2 ¼ 43 Gy) 188 Os (stable, primordial) 143
Pb (stable) ’ 232Th (T1/2 ¼ 14.0 Gy) 207 Pb (stable) ’ 235U (T1/2 ¼ 0.704 Gy) 206 Pb (stable) ’ 238U (T1/2 ¼ 4.47 Gy) 204 Pb (stable, primordial) 10 Be (cosmogenic, T1/2 ¼ 1.5 My) 9 Be (stable, primordial) 208
Hydrothermal solutions at mid-ocean ridges (minor) Today: industrial Pb
10
Be: atmospheric precipitation by rain 9 Be: erosion of the continental crust
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80–200 y (Pacific)
1000 atoms/g
250 y (Atlantic)
0.25 pg g1
600 y (Pacific)
LONG-TERM TRACER CHANGES
by radioactive decay of the parent isotope of at least one isotope in the tracer’s sources or cosmogenic production. The elements currently in use for paleooceanography are given in Table 1. As apparent from the properties listed in Table 1, ocean chemists have a variety of tracers at hand, covering a range of residence times and chemical behaviors. Those tracers varying due to radioactive decay have distinct isotopic compositions in their various source materials (Table 2). This makes them particularly useful both as water mass tracers, and to reconstruct the flux from these various sources into the oceans. It may be surprising to find the cosmogenic nuclide 10Be in this list of otherwise radiogenic tracers. The reason is that Be behaves very similarly to the other tracers in that the ratio 10Be/9Be is distinct in different water masses. Given that 10Be is the only tracer of which the flux into the oceans is known, t can be calculated precisely from its water column concentration. Further, the continent-derived isotope 9Be is the only tracer of which the flux into the oceans can be calculated from the 10Be/9Be ratio. Examples of the isotopes of Nd and Be as water mass labels are shown in Figure 1a and b. The isotope variations of Nd are so small that the 143 Nd/144Nd ratio is reported normalized to a ratio typically found in chondritic meteorites (‘CHUR’): 143
eNd ¼
Nd=144 Ndsample
143 Nd=144 Nd
! 1
104
CHUR
The salinity contours in Figure 1 define water masses, such as North Atlantic Deep Water (NADW), Table 2
Isotope ratios of source materials
Isotope ratio
Pacific mid-ocean ridges
Average upper continental crust
Cosmic dust
87
0.7028 0.125 (abyssal peridotites) 0.5132 þ 10 þ 20 0.2834 18.5 15.5 38.0
0.72 1.26
N/A 0.126
Sr/86Sr Os/188Os
187
143
Nd/144Nd
eNd 176 Hf/177Hf eNd 206 Pb/204Pb 207 Pb/204Pb 208 Pb/204Pb
0.5121 11.4 10 0.2825 19.3 15.7 39.1
N/A N/A
143
Nd=144 Ndsample
143 Nd=144 Nd
CHUR
1
* 10
4
176 eHf ¼
Antarctic Intermediate Water (AAIW), and Antarctic Bottom Water (AABW). Note that eNd is –13.5 in NADW, and 10Be/9Be is c. 0.5 107. In the southern circumpolar water eNd is –9, and 10Be/9Be is 1 107. 10Be/9Be, and in particular eNd, mimic the shape of salinity. Incorporation of these tracers into the sediment at a given location potentially provides information on the distribution and mixing of water masses at this location back through time. The schematic global distribution of deep-water isotope ratios of all tracers discussed here is shown in Figure 4A–F. Note that the variability decreases with increasing residence time. 87Sr/86Sr is perfectly homogenized (Figure 2A). The only location worldwide at which a different Sr isotope ratio has been measured in sea water is the restricted Baltic Sea, where riverine dilution halves the open-ocean salinity and leads to a distinct 87Sr/86Sr only just detectable by modern analytical methods. 187Os/188Os, with an estimated t of 8000–40 000 y, shows only a minute difference between the Atlantic and the other oceans (Figure 2B) show clear gradients between Atlantic and Pacific deep water. This is because the Atlantic receives the highest flux of continental erosion products (aeolian dust, river particulate matter, river dissolved matter) per unit open-ocean area. Furthermore, all this material is derived from old continental crust with an isotope composition distinct from younger rocks (Table 2). Labrador Sea water, for example, receives erosion products from Archean cratons with a unique isotope composition. In contrast, the Pacific receives most of its tracer input from the surrounding volcanic arcs, which have isotope compositions different from the continental crust surrounding the Atlantic. The Indian Ocean has ratios intermediate between the Atlantic and the Pacific for all of these tracers. Whether this is due to mixing of Atlantic water masses (advected through the circumpolar current) and Pacific water (advected via the Indonesian throughflow), or due to internal sources unique to the Indian Ocean is currently not known. 10Be/9Be ratios are lower in the Atlantic because the North Atlantic receives a higher flux of terrigenous 9Be. This keeps the 9Be concentration uniform worldwide, whereas 10Be increases along the advective flow path as expected from a nutrient-type tracer.
N/A
eNd and eHf are 143Nd/144Nd and 176Hf/177Hf ratios, respectively, normalized to a chondritic value CHUR. 143Nd/144NDCHUR ¼ 0.512638; 176Hf/177HfCHUR ¼ 0.282772; (N/A)L: Not Available. eNd ¼
457
Hf=177 Hfsample
176 Hf=177 Hf
CHUR
1
* 10
4
Materials and Methods used in Long-term Tracer Studies It is important that sedimentary materials chosen for long-term tracer studies are true chemical or biogenic precipitates formed in the water column. Contamination by terrestrial detrital material (fine clays from
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458
LONG-TERM TRACER CHANGES Nd Station 315 61°S/62°W _14 _10 _ 6
Station 271 49°S/33°W _6 _12
Station 302 33°S/42°W _6 _12
Station 30 Station 11 36°N/62°W 52°N/47°W _12 _16 _8 _20
Station 63 7°N/40°W _15 _11
0 34
34 5
3 4 .7
5
34. 45
34.50
34 .5 5 34 .6 0
34 .3 5
34 .9 5
34. 85
0
2000
35
0 3 4 .8 .85 34
.0
0
34 .7
3000
5
Depth (m)
.2
34 .3 0
.6
1000
34 .9 5
34 3 4..5 5 60
34 .85
4000 5000 60°S
40°S
20°S
0° Latitude
(A) 10
20°N
0.8
60°N
_
Be/ 9 Be (x10 7) Station 64 Station 22XC Station 8 34°N/63°W 41°N/63°W 45°N/41°W
Station 114 24°S/38°W 0.0
40°N
1.6
0.0
0.4
0.4
0.8
0.4
0.8
0 5
34.45
.6
34.50
34 .5 5 3 4 .6
5
34 .3 5
34 .9 5
34.85
0
2000
0
0
34 .3 0 3 4 .7
35
0 3 4 .8 5 3 4 .8
.0
0
34 .7
3000
5
Depth (m)
.2
34 .9 5
34
1000
34
3 4 .9
34 3 4. 5 5 .6 0
34 .85
4000 5000 60°S
40°S
20°S
0°
20°N
40°N
60°N
Latitude
(B)
Figure 1b (A) Salinity contours of Atlantic sea water with superimposed dissolved Nd isotope compositions. Stippled line gives the typical composition of NADW (eNd ¼ 13). Note the pronounced tongue of NADW with intermediate salinities and eNd of 13 spreading south, and the tongues of AABW and AAIW with lower salinities and eNd of 9 spreading north. (Reprinted with permission from von Blanckenburg F (1999) Tracing past ocean circulation? Science 286: 1862–1863. Copyright 1999 American Association for the Advancement of Science.) (B) Salinity contours of Atlantic sea water with superimposed dissolved Be isotope compositions. Stippled line gives the typical composition of NADW (10Be/9Be ¼ 0.6 107). Note the pronounced tongue of AAIW with lower salinities and 10Be/9Be of 1 107 spreading north. (Data reproduced with permission from Ku et al., 1990; Xu, 1994; Measures et al., 1996).
aeolian or hemipelagic sources) and material affected by chemical alteration through diagenetic processes has to be avoided. The isotopes of Sr are usually measured on carbonates or barite, while those of Be, Nd, or Hf are extracted from chemical sediments, such as ferromanganese (Fe-Mn) crusts, manganese nodules, the authigenic phase of deep-sea sediments, or marine phosphorites, argillites, and glauconites.
Results Sr
Sr isotopes represent the best-studied long-term tracer, as well-dated carbonate sequences are readily available, the extraction and analysis is simple, and the long residence time ensures worldwide homogenization. Therefore, a single isotope evolution curve has
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459
LONG-TERM TRACER CHANGES
87
86
Hf
Sr/ Sr
+8
+1
+6
+1
+5 0.70918
0.70918
+2 +3
0.70918
+3
+4
(A)
+6
(D)
187
Os/188Os
206
Pb/
204
Pb
18.8 18.7
1.06
19.2 18
.8
18.6
19
1.03
1.03
.1
18.8
1.03 18
.9
19.
0
18.8
18 .8
(E)
(B) Nd
10
_ _ 8 13
x10 _7
_7
_7
10
_7
10
_7
0.9x10
8x
1x10
0.
_8
5x
_ 12 _9
1.3
_ 13
_ 6
_7
0.5x10
3x 10 _ 7
0.
_9
1.
_7
_ 4
10
_ 6
1x
_ 5
_20
Be/ 9Be
_1 0
_7
1x10
(C)
(F)
Figure 2A–F (A) Map of modern 87Sr/86Sr ratios in sea water (relative to 0.710248 for the Sr isotope standard SRM 987 (McArthur, 1994). Note that the long t of Sr (several My) allows for perfect homogenization and uniform isotope ratios in all basins. (B) 187 Os/188Os isotope ratios in modern sea water as measured on the surface of hydrogenous Fe-Mn crusts (Burton et al., 1999b). t of a few tens of thousands of years just allows for a small difference between the North Atlantic (with its rich input of old continental material) and the other oceans. (C) Dissolved Nd isotope compositions in modern deep sea water as measured in Mn nodules (reproduced with permission from Albare`de and Goldstein, 1992), adjusted for more recent measurements of Nd in deep sea water (see references in von Blanckenburg, 1999). A t of c. 1000 y allows for distinct gradients between basins, depending on the age and Sm/Nd ratio of their surrounding continental erosion sources. (D) eHf measured in the surface layer of hydrogenetic Fe-Mn crusts and nodules (Albare`de et al., 1998). Similar to Nd, distinct gradients exist between basins. (E) Pre-anthropogenic 206Pb/204Pb ratios measured in the surface layer of hydrogenetic Fe-Mn crusts and nodules. (Abouchami and Goldstein, 1995. Reprinted from Geochimica et Cosmochimica Acta, 60, von Blanckenburg F, O’Nions RK, Hein JR, Distribution and sources of pre-anthropogenic lead isotopes in deep ocean water as derived from Fe-Mn crust, 4957–4963, Copyright (1996), with permission from Elsevier Science.). Pre-anthropogenic Pb cannot be measured in modern sea water which is dominated by industrial Pb. The short t (40–200 y.) results in distinct signatures between and within basins. (F) 10Be/9Be ratios in modern sea water from both the dissolved phase and the surface layer of hydrogenetic Fe-Mn crusts. (Reprinted from Earth and Planetary Letters, 141, von Blanckenburg F, O’Nions RK, Belshaw NS, Gibb A, Hein JR, Global distribution of Beryllium isotopes in deep ocean water as derived from Fe-Mn crusts, 213–226, Copyright (1996) with permission from Elsevier Science.) 10Be is a cosmogenic nuclide that enters the ocean by precipitation from the atmosphere, whereas 9Be is stable and enters the oceans by erosion. The global deep-water t of 10Be is 600 y, allowing for preservation of distinct gradients between the basins.
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460
LONG-TERM TRACER CHANGES
emerged dating back through the entire Phanerozoic (Figure 3), and at a much higher resolution for the Cenozoic (Figure 4). The curve was time-calibrated using biostratigraphy and, in some cases, geochron-
ology of intercalated ash layers. This makes Sr isotopes a very useful tool for stratigraphy. Ages of high confidence can be determined for those periods where 87 Sr/86Sr underwent strong changes. For some periods
87Sr/ 86Sr
0.710
0.708
0.706
Q
Tertiary
Cretaceous Jurassic
0
Triassic
Permian Carboniferous Devonian Silurian Ordovician
200
Cambrian
400
600
Age (Ma) Figure 3 87Sr/86Sr variations for the Phanerozoicum based on samples of brachiopods, belemnites, conodonts, foraminifera, and samples of micritic matrix. (Adapted from Chemical Geology, 161, Veizer J, Ala D, Azmy K et al., 87Sr/86Sr, d13C, d13O evolution of Phanerozoic sea water, 59–88, Copyright (1999), with permission from Elsevier Science.)
0.7092
87
Sr/86Sr
0.7088
0.7084
0.7080
0.7076 0
10
20
30
40
50
60
70
Age (Ma) Figure 4 Cenozoic evolution of marine 87Sr/86Sr, based primarily on analyses of foraminifera from all oceans. Data sources are as in McArthur (1994).
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LONG-TERM TRACER CHANGES
unique ages cannot be assigned. For example, the slope in 87Sr/86Sr in the Early Tertiary period is too low to allow for high-resolution stratigraphy. The changes in the 87Sr/86Sr ratio are controlled by several processes. These are (1) the mid-ocean ridge flux, which is in turn controlled by the spreading rates of the seafloor; (2) the rate of chemical weathering, in particular that of feldspar and calcite; (3) the areal extent of the continents above sea level; (4) changes in the carbonate compensation depth (CCD). Seawater Sr isotope ratios vary within a small range over time. This is because of the long and efficient mixing of Sr, and also because dissolution of continental carbonate buffers seawater Sr isotope compositions within a narrow range. Some of the trends visible in Figure 3 are a relatively slow decrease in 87Sr/86Sr ratios from the Cambrian to the Jurassic period, upon which are superimposed a number of relatively large fluctuations. The sharp decline and following rise at the Permian/Triassic boundary are spectacular, and are thought to reflect either an extreme climate and weathering change, or the sudden mixing of a
461
previously stratified ocean. A second main trend is a relatively rapid increase in 87Sr/86Sr ratios from the Jurassic to the present, upon which a number of relatively small fluctuations are superimposed. Much discussion has been stimulated by the strong Cenozoic increase in 87Sr/86Sr (Figure 4) that has been linked by some workers to the uplift of the Himalayas and the ensuing delivery of high 87Sr/86Sr by Himalayan rivers. This view was challenged by the recent observation that 187Os/188Os (Figure 5), showing a similar and simultaneous increase, cannot be attributed to the dissolved flux draining the rising Himalayas. Therefore there must be a different cause for the rise in Sr too, possibly a worldwide increase in weathering rate. Similar attention was focused on the pronounced rise over the past 2.5 My. One possibility is that the latter can be explained simply by changes in sea level during the glaciations. However, calculations have shown that sea level variations of 200–300 m would be required – far in excess of the c. 100–150 m of change believed to have taken place during the Quarternary period. Therefore a much more plausible explanation is a change in the
1.0 0.9
0.8
0.6
187
Os/
188
Os
0.7
0.5
0.4
0.3 0.2 0.1 0
10
20
30
40
50
60
70
80
Age (Ma) Figure 5 Marine 187Os/188Os record for the past 80 My in all oceans from H2O2-leached metalliferous and hydrogenetic sediments. Note that the pronounced excursion to low ratios at the K/T boundary (65 My) is explained by a meteorite impact. (Reprinted from Geochimica et Cosmochimica Acta, 63, Pegram WJ, Turekian KK. The osmium isotopic composition change of Cenozoic sea water as inferred from a deep-sea core corrected for meteoritic contributions, 4053–4088, Copyright (1999), with permission from Elsevier Science.)
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462
LONG-TERM TRACER CHANGES
terrigenous Sr input supplied by rivers to the oceans. Either a change in the dissolved Sr flux may have occurred or a change in the isotopic composition of rivers, or a combination of both. The increased erosion and availability of weatherable mineral surfaces resulting from the build up of glaciers during the northern hemisphere glaciation might have provided these changes. Os
Because of the exceedingly small differences in Os/188Os between the different ocean basins, Os has a potential similar to Sr to serve as an isotope
187
stratigraphic tool. It is particularly valuable for carbonate-poor pelagic clays and metalliferous sediments, from which Os is extracted by leaching techniques. The Os isotope curve (Figure 5) bears many similarities to the Sr isotope curve (Figure 4), with the exception of three excursions to low 187 Os/188Os ratios. The spectacular drop at the Cretaceous/Tertiary boundary is probably of meteorite impact origin. The second, mid-Paleocene decrease and also the slow recovery following the impact can be explained by exposure of coastal sediments imprinted by the meteoritic Os from the K/ T boundary. The Eocene-Oligocene excursion (B33 Ma) has been explained by an increased supply of
_3 Pacific
_5
_7
Nd (+)
Indian
_9 NW Atlantic
_ 11
_ 13 0
10
20
30
40
50
60
Age (Ma) Figure 6 Nd isotope variations in Cenozoic sea water based on the analyses of hydrogenetic Fe-Mn crusts. Note that despite the assumed t of more than 1000 y the oceans have kept their Nd isotope provinciality observed today throughout the last 50 My. (Reprinted from Earth and Planetary Science Letters, I55, O’Nions RK, Frank M, von Blanckenburg F, Ling HF. Secular variations of Nd and Pb isotopes in ferromanganese crusts from the Atlantic, Indian, and Pacific Oceans, 15–28, Copyright (1998) with permission from Elsevier Science.)
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LONG-TERM TRACER CHANGES
nonradiogenic Os from peridotite weathering (Table 2). As is the case for 87Sr/86Sr, there is a strong increase of 187Os/188Os towards more radiogenic values over the past B14 My. As stated above, this has been linked to the rise of the Himalayas, but recent analyses of Himalayan river waters for Os isotope compositions do not support this view. A more likely possibility is the weathering of ancient crystalline terranes exposed by physical erosion, or the weathering of black shales. These organic-rich sediments have a high Re/Os ratio. Therefore old black shales have the potential to supply Os with a very high 187 Os/188Os ratio to sea water.
19.0
back as 55 Ma. It is thought that despite the large isotopic variability in source materials, the t of Nd is sufficiently long to allow for efficient intra-basin homogenization. This produces the basins’ characteristic Nd isotope blend. Significant variations in eNd are mainly observed for the past 5 My. In the Pacific a decrease in eNd over the past 5 My might be due to an increased flow of AABW (with low eNd, Figure 2C), a rearrangement of the thermohaline circulation following the opening of the Indonesian throughways for exchange of thermocline waters, or an increase in dust input. The strong decrease in north-west Atlantic eNd over the past 3–4 My has been linked to a strengthening of NADW production following closure of the Panama gateway (suppressing northward flow of AABW and AAIW high in eNd ). However, a pronounced decrease of eNd in a shallow ferromanganese crust off Florida has ocurred as early as 8–5 Ma. This has been ascribed to a decreasing inflow of Pacific water through the narrowing Panama gateway. Thus, a change in the amount and style of weathering associated with the onset of northern hemisphere glaciation at 3 Ma is a more likely explanation for the Pleistocene decrease in eNd . In particular the Labrador Sea, a major source of NADW, is surrounded by ancient rocks with eNd as low as 40 and is supplying deep water with eNd of 20 to NADW (Figure 2C). An increase in weathering of this component has the potential to drive the Nd in NADW towards lower compositions.
18.9
Pb
Nd
Nd isotopes, analyzed with low-time resolution in Fe-Mn crusts and given as eNd units, are presented in Figure 6. The most outstanding feature is that the provinciality, observed in eNd of the modern oceans (Figure 2C), has been a feature prevailing as far
NW Atlantic Indian
19.2
Pacific
Pb/ 204Pb
19.1
206
463
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18.7
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18.5 0
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Age (My) Figure 7 Pb isotope variations in Cenozoic sea water based on the analyses of hydrogenetic Fe-Mn crusts. Because of the short t of Pb the oceans have maintained distinct isotope signals throughout the past 50 My. There is more intra-basin variability with time because the short t does not allow such efficient lateral homogenization within basins as is the case for Nd. Therefore, local changes in erosion are much more visible in Pb isotope variations. (Reprinted from Geochimica et Cosmochimica Acta, 63, Frank M, O’Nions RK, Hein JR, Banaker VK, 1689–1708, Copyright (1999) with permission from Elsevier Science.)
No information can be obtained on natural Pb from modern sea water, because of the strong contamination by industrial Pb. Therefore, the pre-anthropogenic Pb distribution has to be obtained from chemical sediments. 206Pb/204Pb time-series, analyzed in Fe-Mn crusts (Figure 7), show patterns of changes that are less clear than those of Nd. Relative differences even within ocean basins are much larger than those observed for Nd. This may be expected from the short residence time of Pb (Table 1), which does not allow for lateral within-basin homogenization to the same degree as Nd. Therefore, local sources dominate the natural Pb budget, and their changes in flux introduce strong isotope variability. For example, in the Indian Ocean a crust located close to the circumpolar current shows a distinctly different history from the more northerly one, experiencing strong changes. The pronounced increase in north-west Atlantic 206Pb/204Pb can be attributed, as eNd , to a change in NADW production, but is more likely due to a change in weathering of the
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LONG-TERM TRACER CHANGES
_7
3.0×10 Pacific Fe _ Mn crusts
_7
Initial
Pacific10Be/9Be
_7
10Be/9Be Meas
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10
_8 _7
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10 20 Depth (mm)
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d
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10 Depth (mm)
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1 (D)
3 5 Age (My)
7
Figure 8 10Be/9Be ratios in Pacific (A, B) and Atlantic (C, D) Fe-Mn crusts (Ling et al., 1997; von Blanckenburg and O’Nions, 1999). The smooth exponential decrease with depth usually observed in the measured data (as fitted following the stippled lines in A, C, where some individual outliers are attributed to alteration) is compatible with radioactive decay of 10Be only (T1/2 ¼ 1.5 My). A kink in the fits to the data is due to changes in growth rates. Changes in the initial ratio would be visible in the form of offsets in the data. At the time resolution of the samples taken (several hundred thousand years) these are not observed. Therefore, the derived growth rates can be used to time-correct the 10Be/9Be ratios for radioactive decay and to calculate initial ratios. The results shown in B and D indicate that the 10Be/9Be ratios in both the Pacific and the Atlantic have been within the range of modern sea water, for the last 7–10 My, despite presumably widely varying erosional input into the ocean.
glaciated areas or a change in the provenance of the erosional products. Be 10
Be/9Be has been analyzed in Fe-Mn crusts as chronometer for the past 10 My. The smooth logarithmic decrease in 10Be/9Be (Figure 8A and C) is compatible with radioactive decay of 10Be. This allows for calculation of growth rates and reconstruction of the initial 10Be/9Be ratios. Quite unexpectedly these initial 10Be/9Be ratios have been within the present-day range of the respective ocean basins through the past 7–10 My. Since the modern ocean basins do display strongly different 10Be/9Be
ratios (Figure 2F), and since a t of c. 600 y favors exchange of Be between basins, the relative constancy of these ratios with time suggests that large-scale changes in deep-water circulation have not taken place within the measured period.
Discussion and Conclusion Changes in the style of weathering have the potential to change the isotope composition of tracers released to the sea. Leaching experiments on fresh mechanically (glacially) weathered rocks have shown a strongly incongruent release (meaning the release of tracers with varying isotope compositions depending
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LONG-TERM TRACER CHANGES
Table 3
465
Incongruent release of isotopes from strongly mechanically weathered continental rocks
Isotope ratio
Isotope change
Original fresh material
Experimental evidence
87
0.725-0.795
Young granitic moraine, Wind River Range, Wyoming, USA
1.5-9.5
Young granitic moraine, Wind River Range, Wyoming, USA
Ammonium acetate leach River water composition HCl leach (Blum and Erel, 1995) HCl leach
Sr/86Sr
187
Os/188Os
eNd
26- 42
206
15.2-22.0
Pb/204Pb
Greenland river bedload Baffin Bay deep-sea sediment Greenland river bedload Baffin Bay deep-sea sediment
on the mode of liberation from rocks) of radiogenic Sr, Os, and Pb, and nonradiogenic Nd (up to 16 eNd units, Table 3). Strongly chemically weathered rocks do not show such an incongruent release. Therefore, some of the long-term changes seen in the isotope evolution of seawater Sr, Os, Nd, and Pb might be attributable to these effects. Certainly the evolution of Sr and Os is only controlled by weathering, and the relative contribution of various sources, such as continental rocks, MORB (Mid Ocean Ridge Basalt), carbonate recycling, peridotite weathering, and cosmic dust. The uniformity between basins makes Sr and Os isotopes reliable stratigraphic tools. 10 Be/9Be appears to be a robust dating tool for the past B10 My, due to its relatively constant observed initial ratio. It remains to be demonstrated whether Nd, Hf, and Pb have real value as tracers of past ocean circulation, as expected from their distinct water mass compositions, or whether their changes back through time merely record local changes in weathering. Clearly the time resolution achievable from Fe-Mn crust studies is not sufficient to answer these questions. Much more insight will be obtained when these radiogenic tracers and Be isotopes are applied to sediments allowing tracer change studies at the resolution of a few thousand years. This will allow more reliable studies on the relationship between climate change, ocean circulation, and continental weathering.
Suggested Reading The topic of radiogenic seawater tracers is too novel to be covered by a single monograph. All information is spread between numerous publications in international journals. Faure (1986) gives a general introduction into radiogenic isotope techniques.
Young terrestrial Fe-Mn coatings Peucker-Ehrenbrink and Blum (1998) HCl leach (own work) HCl and HBr leach (own work)
Analytical methods are summarized in a monograph by Potts (1987). Broecker and Peng (1982) provide a much-cited introduction into the topic of tracers in the sea. Ferromanganese crusts have recently been summarised by Hein et al. (1999). McArthur (1994) has reviewed the material suitable for Sr isotope analysis in carbonates, covering all ages of deposits from recent to the Precambrian. A summary of the suitability of marine clay minerals for isotope analyses is given by Stille et al. (1992). A brief summary on radiogenic seawater tracers, containing useful cross-references, has been published by the author (von Blanckenburg, 1999).
See also Authigenic Deposits. Carbon Cycle. Cenozoic Climate – Oxygen Isotope Evidence. Cenozoic Oceans – Carbon Cycle Models. Cosmogenic Isotopes. Large Marine Ecosystems. Mid-Ocean Ridge Geochemistry and Petrology. Rare Earth Elements and their Isotopes in the Ocean. Redfield Ratio. River Inputs. Sediment Chronologies. Stable Carbon Isotope Variations in the Ocean. Uranium-Thorium Series Isotopes in Ocean Profiles. Water Types and Water Masses.
Further Reading Broecker WS and Peng TH (1982) Tracers in the Sea. Palisades: Lamont-Doherty Geological Observatory. Faure G (1986) Principles of Isotope Geology. John Wiley & Sons. Hein JR, Koschinsky A, Bau M, Manheim FT, Kang JK, and Roberts L (1999) Cobalt-rich ferromanganese crusts in the Pacific. In: Cronan DS (ed.) Handbook of Marine Mineral Deposits, pp. 239--279. Boca Raton: CRC Press.
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McArthur JM (1994) Recent trends in strontium isotope stratigraphy. Terra Nova 6: 331--358. Potts PJ (1987) A Handbook of Silicate Rock Analysis. Blackie. Stille P, Chaudhuri S, Kharaka YK, and Clauer N (1992) Neodymium, strontium, oxygen and hydrogen isotope
compositions of waters in present and past oceans: a review. In: Clauer N and Chaudhuri S (eds.) Isotopic Signatures and Sedimentary Rocks, p. 555. Berlin: Springer Verlag. von Blanckenburg F (1999) Tracing past ocean circulation? Science 286: 1862--1863.
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MACROBENTHOS J. D. Gage, Scottish Association for Marine Science, Oban, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1505–1515, & 2001, Elsevier Ltd.
Introduction The macrobenthos is a size-based category that is the most taxonomically diverse section of the benthos. Only in shallow water does the macrobenthos include both plants and animals. Here, attached macrophytes (including various algae and green vascular plants) may make up a large part of the benthic biomass in coastal areas. Meadows of sea grass (which root into and stabilize coarser sediments) and forests of macroalgae (usually attached to hard bottoms) provide habitat for smaller plant and animal species. In warm, shallow water, stony corals, which flourish as a consequence of symbiotic algae living in their tissues, overgrow large areas and these reefs provide a biogenic habitat for a wealth of other species, and their broken-down skeletons provide much of the sediment in adjacent areas. But the importance of such areas declines rapidly with declining potential for photosynthesis as light quickly vanishes with increasing depth, and the macrobenthos becomes the exclusive domain of heterotrophic life (fueled by breakdown of complex organic material) in soft sediments. Only at deep-sea hydrothermal vents is there an exception to this. These support lush concentrations of benthic biomass that relies not on photosynthetic production ultimately derived from the surface but on the activity of chemoautotrophic bacteria exploiting the emissions of reduced sulfur-containing inorganic compounds. The animal macrobenthos may be attached or may be able to move over hard surfaces provided by exposed bedrock, or may use as habitat the much larger and quantitatively important areas covered by soft sediment. Areas of rock, exposed as a consequence of water currents and turbulence (or, in the case of the newly formed sea floor at the spreading centers along the mid-ocean ridges, rock that has not had time to become covered in sediment settling from above) provide habitat for epibenthos, or epifauna. Even if epibenthos, both plant and animal, looks conspicuous between the tides (and is certainly important
with respect to fouling of colonizers of submerged hard surfaces made by man, such as ships and jetties), it is only a vanishingly small proportion of the huge area of the benthic habitat covering more than half the globe that is not covered by soft sediment. The term infauna has been used to categorize the organisms inhabiting soft sediment. But many epifauna, such as sea stars, are motile and can forage over the surface of sediments. The activities of the animals of the so-called infauna are usually focused on the sediment–water interface where their detrital food is concentrated, but this should not be taken to imply that sediment fauna is always burrowed out of sight. None the less, some species, particularly among larger crustacea, are capable of burrowing even a meter or more deep into the sediment. There are also many species closely related to epifaunal groups of hard substrata, such as sponges and Cnidaria (including sea pens, soft and stony corals and sea anemones), that anchor into the sediment for a sedentary life style, catching small particles from the bed flow.
Global Pattern in Macrobenthic Biomass Extensive Russian sampling after World War II established the precipitous decline in benthic biomass with increasing depth into the abyss. This is caused largely by mid-water consumption of particles escaping from the euphotic zone. Globally, the amount leaving the euphotic zone should be equivalent to the so-called ‘new production’ of roughly 3.4–4.7 109 tonnes C y 1, about 10% of total surface primary production. But this is distributed very unevenly. Perhaps 25–60% is exported in shallow seas, while in the deep ocean only 1–10% reaches the bottom; this is also influenced by latitude-related differences in depth of the mixed layer and intensity of seasonality in the upper ocean. Figure 1 shows how influences such as upwelling and inshore surface productivity will also affect local values of benthic biomass. Overall, these range from highs of more than 500 g m 2 in shallow, productive waters just tens of meters deep, to less than 0.05 g m 2 (equivalent to 2 mg C m2) on the abyssal plains. Trenches are deeper still but, by acting as sumps for material washed in from nearby island arcs and land mass, can support higher than expected biomass.
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Upw elli
ng
Euphotic zone Intertidal zone
Co
ast
al
Oce
Biomass
anic
Euphotic Coast shelf
Aphotic Abyss
Slope
The lower size limit of macrobenthos was determined by Mare as that part retained in a 1 mm sieve, but later became reduced downwards to just 0.5 mm as it was realized not only that small, juvenile sizes were being lost in numbers but also that smaller species of groups already being sampled and taken as part of the macrofauna were not always retained adequately. Mare recognized that the limit might depend on habitat and deep-sea benthic biologists found they had to use even finer-meshed screens to collect the same sorts of animals that characteristically make up the macrobenthos in shallow waters. In the 1970s, Hessler found that he had to use a 297 mm mesh sieve to catch sufficient animal macrobenthos for study in box cores from the abyssal central North Pacific (a very oligotrophic area – nutrient-poor and therefore thin in plankton). Sieves with meshes of just 250 mm are now standard in recent large European studies on the deep-sea macrobenthos.
Depth Figure 1 Conceptual model of food availability and benthic biomass in relation to depth. Upwelling areas provide nutrients for enhanced coastal productivity. The ‘coastal’ curve refers to shelf areas supporting high productivity (usually wider shelves with land inputs, e.g., rivers) compared to the ‘oceanic’ curve where oceanic effects prevail (usually narrow shelves with little land input). The mismatch in the intertidal between biomass and high food availability is explained by the co-occurrence with the latter of high hydrodynamic disturbance by waves and currents. (Modified from Pearson and Rosenberg (1987).)
History and Size Limits of Macrobenthos The term macrobenthos dates from the early 1940s when Molly Mare published a study of an area of coastal soft sediment off Plymouth, England. In recognizing that the benthic ecosystem is fueled by a detrital rain of particles derived from photosynthetic production by macrophytes or phytoplankton, she identified the potential importance of the smaller size classes of metazoan and single-celled organism, right down to bacteria, in the decomposition cycle and food chains in the sediment. She differentiated the benthos into several subcategories based on size, or biomass. Before this time a distinct category for macrobenthos was unnecessary because the only part of the benthos generally thought to be worth studying was the animal life large enough to be eaten by fish. We now differentiate these from the very small metazoans and single-celled algae that Mare named meiobenthos. Another category is the hyperbenthos – small metazoan life that can swim off the bottom and form a distinct community in the benthic boundary layer.
Sources of Food and Feeding Types Patterns in feeding of macrobenthos have often been used to distinguish ecological zones. Although exact definition of feeding category for individual organisms has been controversial, the simplest classification is into suspension and deposit feeders, carnivores, and herbivores. More detailed categorization has proved difficult because of overlap and behavioral flexibility. Although most macrobenthos feed on detrital particles settling from the water column, such as feces, molts, and dead bodies of plankton, this passive sinking is augmented by currents that may resuspend particles periodically from the bottom. Macrobenthos may gather these particles either by catching them from bottom flow or by ingesting the sediment itself as deposit feeders, either in bulk or more selectively for the most nutritious particles. Where currents vary periodically, some animals can feed on both suspended and settled particles by simply changing the way they use their feeding appendages. Just as the particles caught by suspension feeders may range from inert floating detritus up to small swimming organisms, deposit feeding shades into predation where the particles encountered include smaller living benthos. Whether macrofauna can utilize dissolved organic matter in the sediment porewaters to any great extent is still unclear. Wildish has provided the most satisfactory classification of macrofaunal feeding types related to environment. This keeps all three categories but separates deposit feeders into surface and burrowing deposit feeders. Each of the five groups is subdivided
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MACROBENTHOS
in terms of motility and also in terms of food-gathering technique, such as use of jaws and particleentangling structures. These may be arranged along an environmental gradient, such as that illustrated in Figure 2, to allow insight into the causal basis of previously described composition of macrofaunal communities. The relation of feeding to small body size in deepsea macrobenthos may be important. Thiel thought that small body size is a result of a balance between limited food and metabolic rate that makes larger size more efficient than smaller, and of the effects of small population size on reproductive success. Being small allows organisms to maintain higher population densities that increase the chance of encountering the opposite sex and so of reproducing and maintaining the population. The extent of faunal miniaturization is still debated and, surprisingly, not readily summarized by simply taking the total bulk of
the sample and dividing by the number of animals present. The exceptions seem be those organisms that have overcome the reproductive problem by being highly motile scavengers and that need also to be large enough to allow them to forage for the large food falls that occur very sporadically on the deep ocean bed. This scavenger community is quite well developed and includes close relatives of typical macrofaunal organisms in shallow water that in the deep sea grow to a relatively enormous size (Figure 3).
Size Spectra If the sizes of all individuals from an area of sediment are measured and plotted as frequencies along a logarithmic size axis, a pattern of peaks shows up corresponding to the micro-, meio- and macrobenthic size classes (Figure 4). This supports practical intuition but does not explain why such peaks occur (no
50 % 0
HMJ Herbivores HDJ
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FSC FSX
Suspension feeders
FST FMC
FDC SDT
Surface deposit feeders
SDC SMX SST BSX BDX
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BMX BDT CDJ
Carnivores Shallow vegetation a
b c
Shallow Offshore Offshore Offshore Offshore sand sandyrock mud gravel mud f g h e d g k l n p r s u i m o t j
CMJ Deep mud
v
x
w
Figure 2 Distribution of functional groups in boreal coastal macrobenthos (compiled from genera listed by N. S. Jones for the Irish Sea) along an environmental gradient of decreasing food availability and water turbulence and increasing depth and sedimentation. Functional groups: H, herbivore; F, suspension feeder; S, surface deposit feeder; B, burrowing deposit feeder; C, carnivore. Motility: M, motile; D, semi-motile; S, sessile. Feeding habit: J, jawed; C, ciliary mechanisms; T, tentaculate; X, other types. In the upper panel, width of line representing each functional group along the gradient indicates proportional composition at that depth. The lower panel gives a diagrammatic representation of typical feeding position of taxa representative of various groups relative to the sediment–water interface along each gradient. Key to taxa: (a) macroalgae; (b) sea urchins, e.g., Echinus (HMJ); (c) limpets, e.g., Patella (HDL); (d) Barnacles, e.g., Balanus (FSX); (e), (f) serpulids, sabellids (FST); (g) epifaunal bivalves, e.g., Mytilus (FSC); (h) brittle stars, e.g., Ophiothrix (FMC); (i) Venus (FSX); (j) Mya (FSC); (k) Cardium (FDC); (l) Tellinba (FDT); (m) Turritella (SMX); (n) Lanice (SST); (o) Abra (SDC); (p) Spio (SST); (r) Amphiura (FDT); (s) Echinocardium (BMX); (t) Ampharete (SST); (u) Maldane (BSX); (v) Glycera (CDJ); (w) Thyasira (BDX); (x) Amphiura (SDT). From Pearson and Rosenberg (1987).
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1 cm Figure 3 Gigantism in a scavenging amphipod (family Lysianassidae), the cosmopolitan deep-sea species Eurythenes gryllus, compared to the size of a typical shallow-water lysiannasid, Orchomene nana (bottom left), a northern European shallow-water species. (Redrawn from Gage and Tyler (1991) and Hayward PJ and Ryland JS (1995) Handbook of the Marine Fauna of North-West Europe. Oxford: Oxford University Press.)
3
_2
Log10 [biomass (mm m )]
6
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0 _2
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Equivalent spherical diameter (µm)
Figure 4 Macrobenthic size spectra measured from an intertidal inlet in Nova Scotia, subtidal Bay of Fundy, and abyssal sediment from the Nares Abyssal Plain south of Bermuda. The median lines (dotted for intertidal and inshore, and double-dashed for abyssal plain) and range (continuous or solid lines) show a coherent pattern with biomass peaks at 1256 and 8192 mm equivalent spherical diameter (ESD). Downwardpointing arrows indicate minimum detectable biomass. Abundance in the bacterial (leftmost) and meiofaunal (middle) peaks averages 5 103 mm3 m 2. The macrofaunal biomass peak is an order of magnitude higher, but shows greater variability. Biomass in the troughs is about 2–3 orders of magnitude less than adjacent peaks. Sediment was a fluid siltclay. (From Schwinghamer P (1985) Observations on sizestructure and pelagic coupling of some shelf and abyssal benthic communities. In: Gibbs PE (ed) Proceedings of the Nineteenth European Marine Biology Symposium, Plymouth, Devon, U.K. 16–21 September 1984, pp. 347–359. Cambridge: Cambridge University Press.)
to microbenthos and bacteria each particle is a little world on which to attach and grow. Warwick provided a complementary explanation that the peaks also reflect size-related adaptation in the way the life history of the organism is optimized to its environment. For example, larvae of macrofauna exploit the trough between macro- and meiofauna to escape from meiofaunal predators, and thereafter quickly grow into the size range of the ‘macrofaunal’ peak. This is generally lower and less defined than the meiofaunal peak, where organism longevity is just a few weeks at most and there is therefore a narrow range in size, while individual macrofauna might grow over several years so that population size distributions are wider. However, subsequent studies have not found that clear peaks in size spectra occur everywhere. In the deep sea, body size miniaturization does not destroy this pattern, even if the trough at 512–1024 mm between meio- and macrofauna may be less than in coastal sediment (Figure 4). It seems more likely that low food supply has become more important than anything else, so that macrofauna, although settling at roughly the same size, simply do not grow anything like as large as similar coastal species, rather than their somehow perceiving the sediment environment differently from typical macrofaunal organisms in shallow water. We cannot therefore reject the idea that size-based differentiation of the benthos occurs; but is it sensible to stick rigidly to the strict size-based divisions that define the macrofauna as only those organisms within a given range of size, or is it better to compare like with like on the basis of higher taxa determining limits rather than size? With the former definition, the lower limit of the macrobenthos will be determined by size at 1.0 or 0.5 mm, even if this excludes smaller specimens belonging to the same higher taxon, or even much lower-level taxa. This assumes that a size-based functional distinction operates that for the purposes of the study (perhaps environmental impact assessment) will be more important in determining variability than taxonomic affinity. The former function-based definition has been referred to as macrofauna sensu stricto, while the latter, taxonomic one as macrofauna sensu lato.
Composition and Succession such peaks occur, for example, in pelagic communities). Schwinghamer thought that these peaks reflect the way the organism perceives its sediment environment: macrofauna as a continuous medium on, or in, which to move and burrow; meiofauna as a series of interstices between sediment particles; while
Macrobenthos characteristically includes a huge range of phyla (the major divisions of the animal kingdom). In fact, most higher-level taxa are marine and benthic, with most of these part of the macrobenthos. Of the 35 or so known phyla (the major divisions of the animal kingdom), 22 are exclusively
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MACROBENTHOS
marine, with 11 restricted to the benthic environment. Virtually every known phylum is represented in the macrobenthos except for one, the Chaetognatha, or arrow worms (arguably found only in the plankton, although one bottom-living genus is known). This contrasts with the land (including freshwater environments) where only 12 phyla are found (there is only one small, obscure phylum of worm-like animals, the Onychophora, known only on land). This reflects the marine origins of life and the much shorter time of occupation for life on land (barely 400 million years), compared with 800 million years since metazoan organisms first appeared in the ancient ocean. Only five metazoan phyla are normally regarded as part of the next size group down, the meiofauna. The proportional representation of major taxa is conservative, and seems to vary little worldwide with depth, latitude, or productivity regime. It is only in stressed soft-sediment environments, such as those with high organic loading and depleted oxygen (often occurring together), where major departures to this pattern are found (Figure 5). Inshore studies on effects of pollution have contributed to a concept whereby such stress leads to a modified macrobenthos with fewer species, and these characterized by opportunist forms, mainly polychaetes. In tracing recovery after pollution events, it is not clear to what extent a predictable succession occurs. The modern consensus is that there is a random component imposed on a facultative succession in which ‘opportunist’ species pioneer colonization and bring about amelioration in sediment conditions. This allows a more diverse set of species that are more highly tuned to particular habitats to become established through progressively deeper and more extensive bioturbation (Figure 6).
How Many Macrobenthic Species are There? Up to a few years ago it was thought that of the 1.4 to 1.8 million or so species recorded on earth there are perhaps only 160 000 or so known marine species, about 10% of the total. A large-scale sampling programme in deep water off the eastern United States has thrown this into doubt. Along a 180 km section of the continental slope at about 2000 m depth, Grassle and Maciolek found 58% of the species – especially among polychaete (bristle) worms and peracarids (small, sandhopper sized crustaceans) – new to science. The curve of the accumulation of species plotted against increasing area sampled showed no sign of tailing off; the steady increment of new (but rare) species encouraging an
471
extrapolation that this will apply over the wider area of the deep ocean. Depths below the shelf edge cover about 90% of the domain of macrobenthic infauna. But an area of only about 0.5 km2 out of the almost 335 106 km2 area below 200 m depth has yet been adequately sampled for macrofauna using grabs or corers. Because of this huge unexplored area, the actual number of marine macrobenthic species present today is unknown. It must be vastly greater than earlier estimates based on shallow seas, and according to Grassle and Maciolek is conservatively greater than one million, and more likely to rise to 10 million as more of the deep sea is sampled. The overwhelming taxonomic challenge of describing these new species, the painstaking work of sorting samples from the sediment, and the difficulty of seabed experimentation have perhaps slowed progress in understanding the deep-water sediment community. In contrast, much more has been achieved in biological knowledge of hydrothermal vents (and to a lesser extent cold seep communities) since their discovery in 1977.
Large-scale Patterns in Macrobenthic Diversity Large-scale patterns, other than that for biomass, remain controversial. On the basis largely of sampling of the continental shelf, Thorson pointed out that species richness of the epifauna, occupying less than 10% of the total area, and maximally developed intertidally, rises steeply from low levels in the ice-scoured shallows in the Arctic to high values in the tropics. In contrast, the sediment macrofauna he referred to as ‘infauna,’ usually found deeper and unaffected by ice and meltwater, show much less change. This lack of a latitudinal gradient is supported in some other studies. However, latitudinal comparisons by Sanders in the 1960s found depressed diversity in shallow boreal macrobenthos stressed by wide seasonal temperature change compared to the tropics. Thorson through there were about four times more epibenthic than infaunal species, the microhabitat complexity and consequently high species diversification of the epibenthic habitat being much less obvious in sediments. The sameness of this habitat regardless of latitude led Thorson to his concept of parallel level-bottom communities related to sediment type. But several recent studies indicate shallow tropical and deep-sea sediments do not follow this pattern, with much more species-rich communities developing, albeit including lots of ‘rare’ species, in these habitats.
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Macrofaunal abundances Central North Pacific (5600 m)
Tagus Abyssal Plain (5000 m)
Rockall Trough (2875 m)
Firth of Lorne (50 m)
Loch Crevan (50 m)
Setubal Canyon, Portugal (3400 m)
Inner Loch Etive (20 m)
Rame Mud (4 m)
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Figure 5 Proportional representation of the major taxonomic groups of macrofauna in differing soft sediment habitats and depths worldwide. The upper diagrams show this in terms of the relative abundance in these groups; the lower ones in terms of the number of species represented. A broadly similar pattern is shown between both sets although Crustacea are often more abundant in the deep sea rather than shallow-water macrobenthos. Representation of Annelida (mostly polychaete worms) shows most obvious variation in relation to organic carbon loading and oxygen, the three samples from the oxygen minimum zone (OMZ) in the Arabian Sea off Oman showing a pattern of increasing dominance by Annelida, and eventually complete loss of all other groups except Crustacea, with increasing oxygen depletion. (Data from Mare (1942); Gage (1972) Community structure of the benthos in Scottish sea-lochs. I. Introduction and species diversity. Marine Biology 14: 281–297, Gage J (1977) Structure of the abyssal macrobenthic community in the Rockall Trough. In: Keegan BF, O’Ceidigh P and Boaden PJS (eds) Biology of Benthic Organisms (11th European Marine Biology Symposium), pp. 247–260. Oxford: Pergamon Press; Levin LA, Gage JD, Martin C and Lamont PA (2000) Macrobenthic community structure within and beneath the oxygen minimum zone, NW Arabian Sea. Deep-Sea Research II 47: 189–226.)
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1 Aerobic Sediment
2 Anaerobic Sediment
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MACROBENTHOS
4 Normal
Transitory
Typical Macrofauna Dominants
Nucula Amphiura Terebellides Rhodine Echinocardium Nephrops
Labidoplax Corbula Goniada Thyasira Pholoe
Chaetozone Anaitides Pectinaria Myriochele Ophiodromus
Polluted
Grossly polluted
Capitella Scolelepis
No macrofauna Surface covered, by fibre blanket
,
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Figure 6 Changes in macrobenthic fauna along an enrichment–disturbance gradient, such as that associated with pollution. The gradient can be replaced by time in tracing recovery along the x-axis to a ‘normal’ community (left) after a severe pollution event (right). Note the shift in body size (reflecting change from quick-growing and fast-turnover ‘pioneer’ species to slower-growing, longer-lived population species) from right to left, as well as the increased depth and extent of bioturbation in the sediment. There is also a concomitant increase in macrobenthic diversity. (From Pearson TH and Rosenbeg R (1978).)
Depth-related patterns in macrobenthic composition have been studied, particularly for larger benthic invertebrates (technically megabenthos, see discussion earlier) and demersal (bottom-living) fish. Rates of macrofaunal turnover, and clinal variation in individual species, correspond to the rate of change in depth. It is highest in upper bathyal and slowest and most subtle in the abyssal. Many changes in species composition can be related to trophic strategies along a gradient in food and hydrodynamic energy (see Figure 2), but changes in the sort and intensity of biological interactions, such as predation, varying with depth, may also be important, as can life-history characteristics (such as the incidence of planktotropic larval development). Recent studies have also established a degree of pressure adaptation during early development that will further limit vertical range. Such ecological processes must be considered in concert with processes at the evolutionary timescale for understanding of zonation patterns. These processes have been summarized using multivariate statistics. While helping in formulating ideas on causal factors, these may obscure the underlying complexity, which is best understood as the sum of the range and adaptation and evolutionary history of individual species. Rex postulated a mod-slope peak in macrobenthic diversity from studies of sled samples taken throughout the Atlantic by Sanders. However, there is high
variability among individual sample values compared and there are conflicting results from other sites worked in the north-eastern Atlantic. That deep-sea macrobenthos has high species diversity seems well founded, but the extent to which this contrasts with shallow water is unclear. In his original study off the north-eastern United States, Sanders showed an impoverished species richness in samples of macrofauna compared to the adjacent slope and rise. Gray has pointed out that on the outer continental shelf off Norway macrofaunal diversity may be comparable to that found in Sanders’ deep-sea samples, and it is considerably higher still off south-eastern Australia. This suggests not only that the inshore shelf off New England is rather poor in species richness but that the North Atlantic as a whole may be atypical, with perhaps historical factors operating there to restrict macrobenthic diversity compared to the southern hemisphere. Such factors may determine the differing response shown in a comparison of deep-sea macrobenthic diversity among sites throughout the Atlantic where the depression at high latitudes is absent south of the equator. The reduced levels at high latitudes may simply reflect Quaternary glaciation so that the deep Norwegian Sea, isolated by shallow sills from deep water to the south, has a much more recently diverged and quite distinct macrofauna from that in the Atlantic.
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Small-scale Pattern In sampling macrobenthos from a ship it is easy to assume that the animals in this apparently homogeneous habitat are randomly distributed, but sample replicates may not always provide good estimates of the population mean and its sampling error. When samples are mapped over the sediment, or variability is analysed in large numbers of replicates, clumped distributions of some kind are commonplace (Figure 7). Nonrandomly even (regular) dispersions have been detected, but only at the centimeter scale, suggesting that they are actively defined by the ambit (such as the area swept by feeding tentacles) of individual animals. To describe rather than just detect such nonrandom spatial pattern has been a challenging task, not least because most macrofauna are not readily visible in seabed photographs and the very analysis of samples by sieving and mud will destroy fine-scale pattern. Spatial pattern is usually envisaged in the horizontal plane because macrobenthic organisms concentrate their activity on the sediment–water interface in feeding, movement, and reproduction. Pattern is a dynamic expression of this and consequently may change through time but marine sediments provide a three-dimensional habitat so that vertical as well as horizontal spatial patterns may occur. The latter may be best developed at the small scale where smaller macrofauna (and meiofauna) are concentrated around irrigatory or feeding burrows of larger species (Figure 8). The problem is that to
No. of individuals
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analyze this pattern it is difficult not also to disrupt the habitat. Yet in order to understand the basis of such pattern it is vital to analyze and map pattern over a range of scales in conjunction with variability in the sediment habitat. Some of the most revealing studies have examined dispersions of individual species over plots measuring tens of meters square. These may reveal the two aspects of pattern, intensity and form. Intensity can relatively easily be measured by the ratio of variance to mean. This will distinguish distributions that are clumped, regular, or not statistically distinguishable from random. The form of pattern is an aspect that classical statistical tests of nonrandomness do not address. Yet a clumped pattern may be very different in form from that shown by another species that shows similar intensity of aggregation. Although a nuisance for the easy interpretation of sample statistics, an understanding of the biological basis of patterns will provide important insight into the processes maintaining macrobenthic communities.
Altered erosion
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10 cm Figure 7 Upper, three-dimensional plots of abundance showing spatial dispersion patterns of two infaunal bivalve molluscs Nucula hartivigiana (A) and Soletellna siliqua (B) in a 9000 m2 area of mid-tide sandflat with no obvious gradients in physicochemical conditions. Both species show quite different spatial patterns. The lower plots show spatial correlograms with significant autocorrelation coefficients (measured as Moran’s I) denoted by filled circles (From Hall et al. (1984). Thrush et al., (1989).)
Anaerobic sediment
Figure 8 Benthic biological activity and seabed sediment structure. The diagram shows some of the ways macrobenthos, in conjunction with other size classes, influence sediment fabric, physicochemical properties and solute fluxes (see also Table 1). (After Meadows PS (1986) Biologica activity and seabed sediment structure. Nature 323: 207.)
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Functional Importance of Macrobenthos Using a grab as a quantitative sampler in the early years of the twentieth century, C. J. J. Petersen hoped to be able to work out from his thousands of samples taken in the North Sea how much food was available to fish such as flounder. Such links were well supported from finds of large numbers of benthic animals in fish guts, even if later work showed that fish are by no means as important predators of macrobenthos as are invertebrates such as sea stars. Petersen also noticed that characteristic and uniform assemblages of macrobenthos were found that could be related to sediment type and that provided him with statistical units that Thorson later used as his descriptive units in his concept of ‘parallel level bottom communities.’ Ecologists debated whether these existed as anything more than assemblages responding to similar conditions (as originally inferred by Petersen) or reflected functional units or ‘biocoenoses’ where biological interactions play an important, if unknown, role. However, the importance of biological interactions in the subtidal community is difficult to address experimentally, and most data are available from intertidal mudflats and sandflats where ecological gradients related to tidal exposure pose additional complexity. Manipulative experiments on the effects of predation by caging small areas show that predators like shore crabs can have big effects on prey densities, while other studies show competitive exclusion between worms with different burrowing styles that can be reflected in clumping patterns. This contrasts with the importance of grazers and predators in preventing dominance by fast-growing competitive superior species on rocky shores. Such biological interaction cascades down through the community – so-called ‘top-down’ control. But, in sediments, effects such as predation are not so marked overall. Perhaps the three-dimensional structure, the uneven distribution of food and irrigatory flows, and the often intense stratification of chemical processes reduce competition. Indirect effects, such an bioturbation, may take the place of competition. By bulk processing large quantities of sediment, large macrofaunal deposit feeders rework the sediment down to the greatest ocean depths and thus exert a major influence on benthic community structure. It has been suggested that this constant process of biogenic disturbance and alteration of the benthic environment by macrofauna (Figure 8, Table 1) encourages high species richness among smaller macrobenthos by reducing them to levels where competition is relaxed, so that more species can coexist. It is also argued that the constantly
475
changing micro-landscape created by other, larger species provides a rich niche variety for macrofauna. It is difficult to see this process operating on the vast abyssal plains where faunal densities, and therefore such biogenic effects, are so low but species richness is high. Grassle’s spatiotemporal mosaic theory sees the deep-sea bed having patchy and ephemeral food resources that create a relatively small, discrete, and widely separated patch structure promoting coexistence. Effects of larger-scale disturbances are more difficult to detect let alone manipulate in experiments with the sediment community. Yet the evidence is that physical disturbance such as that caused by storm-driven sediment scour and resuspension may have an important effect on assemblage structure and species richness on the exposed continental shelf and margin. The expectation that, just as on an exposed sandy shore, only a relatively small suite of species will be able to adapt to such conditions is confirmed in the deep sea on the continental rise off Nova Scotia, where benthic storms occur with relatively high frequency. Benthic storms may occasionally occur on the abyssal plains, so it should not be assumed that biogenic structure is simply longer-lasting there because it takes so long to be covered by the very low rate of natural sedimentation. Perhaps the most important determinant of the macrobenthic assemblage, or community, is the larval stage usually dispersed in the water column. Larvae can test the substratum and swim off until they find conditions suitable for settlement and metamorphosis. On rock this may involve a series of precise cues that can include presence of their own or other species. Less is known about settlement of infaunal species, but it is thought that positive cues such as microtopography may be much less important in sediment dwellers, while negative cues such as the presence of other species or unattractive sediment are more important. Nevertheless, a community will still be very largely constrained by supply of propagules. In a coastal area the access of larvae supply from adjacent breeding populations may be constrained by coastal topography and currents, not to mention barriers formed by features such as estuaries. The patch structure in the deep sea is maintained by water-borne dispersal stages with the resulting metapopulations spatially unautocorrelated (presence of an organism not dependent on other occurrences). In deep water the openness of the system may mean that the sediment is exposed to a much larger pool of species, even if they are at very low densities as larvae. In the tropics a similar effect results from the greater incidence of planktotrophy (larvae feeding in the plankton), when the longer
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Table 1
Direct and indirect effects of macrofauna on soft sediments and their ecological consequences
larval life will ensure wider dispersal than that of the nonfeeding larvae prevalent in cooler waters. This may help to explain why so many, mostly rare, species can coexist in both environments.
Importance of Macrobenthos in Environmental Assessment Because the benthic community (unlike fish or plankton) is stationary or at best slow-moving over a small area of bottom, it is useful in monitoring environmental change caused by eutrophication and chemical contamination. Macrobenthos studies have defined the generic effects of such sources of stress by changing representation of major taxa, reduction in diversity, and increasing numerical dominance by small-sized opportunist species causing a downward shift in size structure. This seems to be accompanied by greater patchiness, reflected by increased variability in species abundances in sample replicates. It is also seen as greater variability in local species diversity caused by greater heterogeneity in species identities. This reflects subtle changes in abundance
and, particularly in the more species-rich communities, changes in presence/absence of rare species that might be detected earlier at less severe levels of disturbance. It is claimed that in bioassessments comparing species richness using samples of macrobenthos rare species should receive greater attention by taking larger samples because they contribute relatively more to diversity than the abundant community dominants. Other workers argue that very many species, especially rare ones, are interchangeable in the way they characterize samples. This question requires investigation of the way stressors impact the community, and whether it is the dominant or the rare species that are most sensitive, and therefore most rewarding for study in detecting impacts. Interpretation of impacts also has to proceed against a background of natural changes in benthic communities caused by little-understood, year-toyear differences in annual recruitment. In establishing a baseline there is a need also to take into account the little-understood effects of bottom trawling on coastal benthos. Such disturbance in parts of the North Sea may date back at least 100 years, and now
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MACROBENTHOS
means that virtually every square meter of bottom is trawled over at least once a year. Such monitoring has in the past entailed costly benthic survey and tedious analysis of samples to species level. Consequently, there has been effort to see whether the effects of stress can be detected at higher taxonomic levels, such as families. Higher taxonomic levels may more closely reflect gradients in contamination than they do abundance of individual species because of the statistical noise generated from natural recruitment variability and from seasonal cycles such as reproduction. This hierarchical structure of macrobenthic response means that, as stress increases, the adaptability of first individual animals, then the species, and then genus, family, and so on, is exceeded so that the stress is manifest at progressively higher taxonomic level. Such new approaches, along with the nascent awareness of conservation of the rich benthic diversity, and with a need for improved environmental impact assessment on the deep continental margin, should ensure a continued active scientific interest in macrobenthos in the years to come.
See also Benthic Boundary Layer Effects. Benthic Foraminifera. Benthic Organisms Overview. Coral Reefs. Deep-Sea Fauna. Demersal Species Fisheries. Fiordic Ecosystems. Grabs for Shelf Benthic Sampling. Meiobenthos. Microphytobenthos. Phytobenthos. Pollution: Effects on Marine Communities. Rocky Shores. Sandy Beaches, Biology of
Further Reading Gage JD and Tyler PA (1991) Deep-sea Biology: A Natural History of Organisms at The Deep-sea Floor. Cambridge: Cambridge University Press.
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Graf G and Rosenberg R (1997) Bioresuspension and biodeposition: a review. Journal of Marine Systems 11: 269--278. Gray JS (1981) The Ecology of Marine Sediments. Cambridge: Cambridge University Press. Hall SJ, Raffaelli D, and Thrush SF (1986) Patchiness and disturbance in shallow water benthic assemblages. In: Gee JHR and Giller PS (eds.) Organization of Communities: Past and Present, pp. 333--375. Oxford: Blackwell. Hall SJ, Raffaelli D, and Thrush SF (1994) Patchiness and disturbances in shallow water benthic assemblages. In: Giller PS, Hildrew HG, and Raffaelli DG (eds.) Aqautic Ecology: Scale, Patterns and Processes, pp. 333--375. Oxford: Blackwell Scientific Publications. Mare MF (1942) A study of a marine benthic community with special reference to the micro-organisms. Journal of the Marine Biological Association of the United Kingdom 25: 517–554. McLusky DS and McIntyre AD (1988) Characteristics of the benthic fauna. In: Postma H and Zijlstra JJ (eds.) Ecosystems of the World 27, Continental Shelves, pp. 131--154. Amsterdam: Elsevier. Pearson TH and Rosenberg R (1978) Macrobenthic succession in relation to organic enrichment and pollution of the marine environment. Oceanography and Marine Biology: an Annual Review 16: 229--311. Pearson TH and Rosenberg R (1987) Feast and famine: structuring factors in marine benthic communities. In: Gee JHR and Giller PS (eds.) Organization of Communities: Past and Present, pp. 373--395. Oxford: Blackwell Scientific Publications. Rex MA (1997) Large-scale patterns of species diversity in the deep-sea benthos. In: Ormond RFG, Gage JD, and Angel MV (eds.) Marine Biodiversity: Patterns and Processes, pp. 94--121. Cambridge: Cambridge University Press. Rhoads DC (1974) Organism–sediment relations on the muddy sea floor. Oceanography and Marine Biology Annual Reviews 12: 263--300. Thorson G (1957) Bottom communities (sublittoral or shallow shelf). In: Hedgepeth JW (ed.) Treatise on Marine Ecology and Paleoecology, pp. 461--534. New York: Geological Society of America. Thrush S (1991) Spatial pattern in soft-bottom communities. Trends in Ecology and Evolution 6: 75--79.
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MAGNETICS F. J. Vine, University of East Anglia, Norwich, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1515–1525, & 2001, Elsevier Ltd.
intensity’), is the tesla (T). Because the magnitude of the Earth’s magnetic field and magnetic anomalies is very small compared to 1 T, they are usually specified in nanoteslas (nT) (1 nT ¼ 109 T). Fortunately, an old geophysical unit called the gamma is equivalent to 1 nT.
Introduction Since World War II it has been possible to measure the variations in the intensity of the Earth’s magnetic field over the oceans from aircraft or ships. In the 1950s the first detailed magnetic survey of an oceanic area, in the north-east Pacific, revealed a remarkable ‘grain’ of linear magnetic anomalies, quite unlike the anomaly pattern observed over the continents. In the 1960s it was realized that these linear anomalies result from a combination of seafloor spreading and reversals of the Earth’s magnetic field. Hence they provide a detailed record of both the evolution of the ocean basins, and the timing of reversals of the Earth’s magnetic field, during the past 160 million years. In addition, because of the dipolar nature of the field and the dominance of ‘fossil’ magnetization in the oceanic crust, the linear anomalies formed at midocean ridge crests, and the anomalies developed over isolated submarine volcanoes, can sometimes yield paleomagnetic information, such as the latitude at which these features were formed.
Units In the SI system the unit of magnetic induction, or flux density (which geophysicists refer to as ‘field
(A)
History of Measurement William Gilbert, one-time physician to Elizabeth I of England, is thought to have been the first person to realize that the form of the Earth’s magnetic field is essentially the same as that about a uniformly magnetized sphere. This is also equivalent to the field about a bar magnet (or magnetic dipole) placed at the center of the Earth, and aligned along the rotational axis (Figure 1). Certainly in terms of the written historical record he was the first person to propose this, in his Latin text De Magnete, published in 1600. Presumably, with the extension of European exploration to more southerly latitudes in the late fifteenth and the sixteenth centuries, mariners had problems with their compasses that Gilbert realized could be explained if the vertical component of the Earth’s magnetic field varies with latitude. Accurate measurements of the direction of the Earth’s magnetic field at London date from Gilbert’s time. Measurement of the strength or intensity of the field, however, was not possible until the early part of the nineteenth century. The equipment used then, and for the following one hundred years or so, included a delicate suspended magnet system and required accurate orientation and leveling before a measurement
N
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Figure 1 The (A) normal (present-day) and (B) reversed states of the dipolar magnetic field of the Earth. Shaded area shows the Earth’s core; heavy arrows indicate directions of the field at the Earth’s surface.
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MAGNETICS
could be made. Measurements on a moving platform such as a ship were extremely difficult, therefore, and with the advent of iron and steel ships became impossible because of the magnetic fields associated with the ships themselves. Meanwhile, measurements on land had revealed that although the Earth’s magnetic field is, to a first approximation, equivalent to that about an axial dipole as envisaged by Gilbert, there is a considerable nondipole component, on average 20% of the dipole field. Moreover, the field is changing in time, albeit slightly and slowly, in both intensity and direction. As rather more than 70% of the Earth’s surface is covered by water and there were few magnetic measurements in these areas, detailed mapping of the field at the surface worldwide was seriously hampered. In 1929 the Carnegie Institution of Washington went to the length of building a wooden research ship, the Carnegie, to map the Earth’s magnetic field in oceanic areas. Similarly the then USSR commissioned a wooden research ship, the Zarya, in 1956. However, by this time new electronic instruments had been developed that were capable of making continuous measurements of the total intensity of the Earth’s magnetic field from aircraft and ships. Among the many projects instigated by the Allies in World War II, to counter the submarine menace, was the Magnetic Airborne Detector (MAD) project. The outcome was the development of the fluxgate magnetometer. To increase the sensitivity of the instrument, much of the Earth’s magnetic field is ‘backed off’ by a solenoid producing a biasing field. Initially it was not possible to produce a constant biasing field and the instruments tended to ‘drift,’ which meant that they were not ideal for scientific purposes. After the war these instruments were redeployed for use in conducting aeromagnetic surveys over land areas, in connection with oil and mineral exploration, and then modified for use from ships. In both instances the detector was housed in a ‘fish’ that was towed, so as to remove it from the magnetic fields associated with the ship or aircraft. In the 1950s the proton-precession magnetometer was developed, which had the advantage of achieving the same or somewhat better sensitivity (about 1 part in 50 000) without the problem of drift. Since 1970 even more sensitive magnetometers have been developed – the optical absorption magnetometers. Although in some ways superseded by these magnetometers based on proton and electron precession, fluxgates are still widely used because they have the advantage of measuring the component of the field directed along the axis of the detector rather than the total ambient field. This property makes them particularly useful in satellites, for example, where they
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are often used as components in orientation systems, at the same time yielding measurements of the Earth’s magnetic field. Measurements of the Earth’s magnetic field in oceanic areas are, therefore, now relatively routine, whether they be from satellites, aircraft, ships, submersibles, or remotely operated or deeply towed vehicles.
Nature of the Earth’s Magnetic Field The differences between the measured magnetic field about the Earth and that predicted for a central and axially aligned dipole are considerable. Best-known is the difference between the magnetic poles – where the field is directed vertically – and the rotational, geographic, poles of the Earth. As a result, for most points on the Earth’s surface there is an angular difference between the directions to true north and to magnetic north. Mariners refer to this as the magnetic variation; scientists refer to it as the magnetic declination. The centered dipole that best fits the observed field predicts a field strength of 30 000 nT around the equator and a maximum value of 60 000 nT at both poles. The actual field departs considerably from this, as can be seen in Figure 2. The intensity and direction of the field, or any of their components (such as magnetic variation) at any one point, vary with time, typically by tens of nanoteslas and a few minutes of arc per year. This is known as the secular variation of the field. The form of the Earth’s magnetic field and its secular variation is thought to derive from the fact that it is generated in the outer, fluid core of the Earth, which is metallic (largely iron and nickel) and hence a good electrical conductor. Convective motions of this fluid conductor carrying electrical currents and interacting with magnetic fields produce a dynamo-like effect and an external magnetic field. The essential axial symmetry of this field is probably determined by the influence of the Coriolis force on the precise nature of the convective motions. Historical, archeomagnetic and paleomagnetic data suggest that although the field at any one time departs considerably from that predicted by a geocentric axial dipole, when averaged over several thousand years, the mean field is very close to that about such a dipole. On even longer, geological, timescales the field intermittently reverses its polarity completely (Figure 1), probably within a period of approximately 5000 years. The length of intervals of a particular polarity varies widely from a few tens of thousands of years to a few tens of millions of years.
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MAGNETICS 80°N
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Figure 2 Total intensity of the Earth’s magnetic field (in nT) at the Earth’s surface (for epoch 1980). (Reproduced from Langel RA, in Geomagnetism Vol. 1, JA Jacobs (ed.) ^ 1987 Academic Press, by permission of the publisher.)
The secular variation and changes in the polarity of the field result from dynamic processes deep in the Earth’s interior. There are also shorter-period variations of the field with time that are of external origin, essentially a result of the interaction of the solar wind with the Earth’s magnetic field. The most relevant of these in the present context is the daily or diurnal variation of the field. This is a smooth variation that is greatest during daylight hours and typically has an amplitude of a few tens of nanoteslas. It can be greater, however, near the magnetic equator and poles. Increased solar activity can produce higher-amplitude and more irregular variations, and intense sunspot activity produces global magnetic storms; high-amplitude, short-period variations during which magnetic surveying has to be discontinued.
Reduction of Magnetic Data Measurements of the Earth’s magnetic field over oceanic areas are corrected for the present-day spatial and time variations of the field described above in order to obtain residual ‘anomalies’ in the field. These are caused by magnetization contrasts at or near the Earth’s surface, i.e., in the upper lithosphere. Such anomalies should therefore yield information on the magnetization and structure of the oceanic crust. For measurements made at or near the Earth’s surface, i.e. from ships, aircraft, or submersibles, the secular and diurnal variation of the field can often be
ignored because these effects are very small compared to the amplitudes of the anomalies being mapped. However, should a very accurate survey be required, secular variation has to be taken into account if surveys made at different times are being combined and the observations should be corrected for diurnal variation using records from nearby land stations or moored buoys. If there are sufficient ‘cross-overs’ during the survey (i.e. repeat measurements at the same point), it may also be possible, indeed preferable, to use these to correct for diurnal variation. It may also be necessary to correct for any magnetic effect of the moving platform itself. If present, this effect will vary according to the direction of travel. For many purposes, however, the above corrections are so small that they can be ignored. The final correction, the removal of the main or ‘regional’ field of the Earth, that is, the field generated in the Earth’s core, must always be applied. In theory this should be simple. The depth to the core–mantle boundary, 2900 km, means that the field originating in the core should have a smooth, long-wavelength variation at or above the Earth’s surface. The magnetization contrasts within a few tens of kilometers of the Earth’s surface will produce anomalies of much shorter wavelength. Between these two source regions any magnetic minerals in the mantle are at a temperature above their Curie temperature and are effectively nonmagnetic. In practice, because of its complexity and because it is changing with time, it
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MAGNETICS
has proved difficult to accurately define the main field of the Earth, i.e. that originating in the core, and the way in which it is changing with time. The first attempt to define a global ‘International Geomagnetic Reference Field’ was made in the 1960s and, although revised and greatly improved at fiveyear intervals since then, its level, if not its gradients, can still be seen to be slightly incorrect for certain oceanic areas. As a result, particularly in the past, the regional field for particular profiles or surveys has been obtained by fitting a smooth long-wavelength curve or surface to the observed data. Once the long-wavelength ‘regional field’ has been removed from magnetic data, the resulting ‘residual’ or ‘total-field’ anomalies are assumed to result from magnetization contrasts within the upper, magnetic, part of the Earth’s lithosphere (Figure 3).
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Figure 3 Magnetic profile recorded by a fluxgate magnetometer between the Cape Verde islands and Dakar, Senegal. The ship’s track close to the islands is shown in the upper part of the diagram. The dashed line indicates the regional field used to calculate total field magnetic anomalies. The highamplitude, short-wavelength anomalies close to the Cape Verde Islands reflect the presence of highly magnetic volcanic rocks at shallow depth. (Reproduced from Heezen BC, et al., in Deep-Sea Research, Vol. 1 ^ 1953 Elsevier Science, by permission of the publisher.)
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Magnetization of Ocean Floor Rocks All minerals, and hence all rocks, exhibit magnetic properties. However, the magnetization that most rock types acquire in the relatively weak magnetic field of the Earth is insufficient to produce significant anomalies in the Earth’s magnetic field, particularly when this is measured at some distance from the rocks, as is typically the case in oceanic areas. For rocks to be capable of producing appreciable anomalies they must contain more than a few percent by volume of ‘ferromagnetic’ minerals, that is, certain oxides and sulfides containing iron, notably magnetite (Fe3O4). Most sediments and ‘acid’ (silicarich) igneous rocks, such as granite, do not meet this criterion. Basic (silica-deficient) igneous rocks such as basalts, and the coarser-grained but chemically equivalent gabbros, and ultrabasic rocks such as peridotite do contain a higher proportion of iron oxides and are capable of producing anomalies. Metamorphic rocks, formed when preexisting rocks are subjected to high temperatures and/or pressures, are typically weakly magnetized except for some formed from basic or ultrabasic igneous rocks. Apart from its sedimentary veneer, the upper part of the oceanic lithosphere consists almost entirely of basic and ultrabasic rocks, i.e. basalts, gabbros, and peridotites. This is a consequence of the way in which it is formed by the process of seafloor spreading. At midocean ridge crests the ultrabasic peridotite of the Earth’s mantle undergoes partial melting, producing basic magma that rises and collects as a magma chamber within oceanic crust. Solidification of such magma chambers ultimately forms the main crustal layer of gabbro, but not before some ultrabasic rocks have formed at the base of the chamber from the accumulation of first formed crystals, and magma has been extruded through near vertical fissures to form pillow basalts on the seafloor. Solidification of the magma in these fissures forms a layer consisting of dikes between the gabbro and the basalts. Thus, because of the rock types present, the oceanic crust and upper mantle are relatively strongly magnetized and capable of producing largeamplitude anomalies in the Earth’s magnetic field, even when measured at sea level. The thickness of the magnetic layer is determined by the depth to the Curie point isotherm. Because of the way in which the oceanic lithosphere is formed, by spreading about ridge crests, this varies from a few kilometers depth within the crust at ridge crests to a depth of approximately 40 km in oceanic lithosphere that is 100 million years, or more, in age. In places the thermal regime associated with seafloor spreading is modified by mantle ‘hot spots.’ As a result there is an
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enhanced degree of partial melting of the mantle and the larger volumes of magma produced extrude on to the seafloor to form seamounts, oceanic islands and, exceptionally, oceanic plateaux. These all involve an appreciable thickening of the oceanic crust, but the rock types involved are all essentially basic and potentially strongly magnetic (Figure 3).
Observed Anomalies The marked contrast in the way in which continental and oceanic crust are formed, and hence in the 135°W
predominant rock types in each setting, gives rise to a striking difference in the character of the total field anomalies developed over the two types of lithosphere. Within the continents the variety of rock types in mountain belts and their juxtaposition by folding and faulting produces magnetization contrasts and anomalies that delineate the general trend of the belt. Areas of igneous activity that include basic igneous rocks are characterized by very largeamplitude and typically short-wavelength anomalies, and sedimentary basins and extensive areas of granite are quiet magnetically. This pattern of
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Figure 4 Linear magnetic anomalies in the north–east Pacific. Areas of positive anomaly are shown in black. Straight lines indicate faults offsetting the anomaly pattern; arrows, the axes of three short ridge lengths in the area – from north to south, the Explorer, Juan de Fuca and Gorda ridges. (Based on Figure 1 of Raff AD and Mason RG in Bull. Geol. Soc. Amer., Vol 72. ^ 1961 Geological Society of America. Reproduced by permission of the publisher.)
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MAGNETICS
anomalies is characteristic of the continental shelves out to the continental slope, the true geological boundary between the continents and the oceans, although, in that the shelves are typically underlain by a great thickness of sediments, they are often magnetically ‘quiet.’ Deep sea areas, that is, those underlain by oceanic lithosphere, are characterized by remarkably linear and parallel anomalies that extend for hundreds of kilometers and are truncated and offset by fracture zones (Figure 4). The fracture zones are formed by the transform faults that offset the crest of the midocean ridge system. Thus the linear magnetic anomalies parallel the ridge crests. The anomalies are remarkable for their linearity, their high amplitude and the steep magnetic gradients that separate highs from lows. Any explanation of them in terms of linear structures and/or lateral variations in rock type within the oceanic crust is extremely improbable. It transpires that they result from a combination of sea floor spreading and reversals of the Earth’s magnetic field. As new oceanic crust and upper mantle form at a ridge crest they acquire a permanent (remanent) magnetization which parallels the ambient direction of the field. If, as spreading occurs, the Earth’s magnetic field reverses, then the ribbon of newly formed oceanic lithosphere along the whole length of the spreading ridge system acquires a remanent magnetisation in the opposite direction. It is these contrasts between normally and reversely magnetized material which produce the high amplitude linear anomalies and the steep gradients between them. Rates of seafloor spreading vary greatly for different ridges and the interval between reversals of the Earth’s magnetic field is also very variable throughout geological time. However, rates of spreading are typically a few tens of millimeters per year and the average polarity interval is about 0.5 million years. Thus typical linear anomalies are 10–20 km in width. Initially, spreading rates could only be reliably determined for the past 3.5 million years. For this period the reversal timescale had been independently determined from measurements of the age and polarity of remanent magnetization of both subaerial lava flows and deep-sea sediments, and it is clearly reproduced in the anomalies recorded across midocean ridge crests (Figure 5). With the dating of older oceanic crust by the international Deep Sea Drilling Program it became possible to deduce spreading rates at earlier times and to calibrate the timescale of reversals of the Earth’s magnetic field implied by the older linear anomalies. In this way the geomagnetic reversal timescale for the past 160 million years has been deduced (Figure 6).
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Figure 5 (A) Schematic crustal model for the Juan de Fuca Ridge, south-west of Vancouver Island. Shaded material in layer 2 is normally magnetized; unshaded material is reversely magnetized. SL ¼ sea level. (B) Part of the summary map of magnetic anomalies recorded over the Juan de Fuca Ridge (Figure 4). (C) Total field magnetic anomaly profile along the line indicated in (B). (D) Computed profile assuming the model and reversal timescale for the past 3.5 million years. (Reproduced from Vine FJ in The History of the Earth’s Crust. RA Phinney (ed.). ^ 1968 Princeton University Press, by permission of the publisher.)
In recording the times at which the Earth’s magnetic field has reversed its polarity, the linear magnetic anomalies also serve as time or growth lines that reveal the evolution of the ocean basins in terms of seafloor spreading (Figure 7). Thus it is possible to accurately reconstruct the most recent phase of continental drift (during the past 180 million years) when a former supercontinent was split up to form the present-day continents and the Atlantic and Indian Oceans. The Pacific Ocean was formed during the
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Figure 6 Geomagnetic polarity timescale for the past 160 million years. (Reproduced from Jones EJW, Marine Geophysics. ^ 1999 John Wiley and Sons, by permission of the publisher.)
same period, and in the process dispersed marginal fragments of the supercontinent around its rim where they are now recognized as ‘suspect terranes.’ In contrast to this very detailed record of spreading and drift for the past 180 million years there is no such record for the previous 96% of geological time, and earlier phases of drift and mountain building have to be deduced from the more complex and fragmentary geological record within the remaining 40% of the Earth’s surface that is covered by continental crust.
Paleomagnetic Information Contained in Oceanic Magnetic Anomalies As a result of the dipolar nature of the Earth’s magnetic field, whereby its inclination to the horizontal varies systematically from the equator to the
poles (Figure 1), anomalies in the total field over a relatively simple and symmetrical feature such as the central, normally magnetized ribbon of crust at a midocean ridge crest typically have an asymmetry (Figure 8). The exceptions occur at the poles and across ridges trending exactly north–south, where the anomaly is a symmetrical high, and over an east– west trending ridge at the Equator, where the anomaly is a symmetrical low. For all other latitudes, and all orientations other than north–south, the degree of asymmetry, or the ‘phase-shift’ of the anomaly, is a function of the latitude and orientation. As a result of spreading, all older linear anomalies, unless formed about a north–south trending, east– west spreading ridge, will now be at a different latitude from that at which they were formed, and the direction of their remanent magnetization will be
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MAGNETICS
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Figure 7 A compilation of the trends of linear magnetic anomalies in (A) the Atlantic and Indian Oceans and (B) the Pacific Ocean. (Reproduced from Jones EJW, Marine Geophysics. ^ 1999 John Wiley and Sons, by permission of the publisher.)
different from the direction of the ambient magnetic field. As a consequence, the asymmetry of the anomaly is different from what one would predict for similarly directed remanence and ambient field. This difference between the observed and predicted phase shift reflects the latitudinal change. Although the interpretation of such data is somewhat ambiguous if the orientation of the anomaly is thought to have changed since the time of formation, it does provide paleolatitude, i.e. paleomagnetic, information for oceanic areas, which is otherwise rather sparse. As with all paleomagnetic data, they can be used to test independently derived models for the ‘absolute’ motion of plates and plate boundaries, such as ridge crests, across the face of the Earth.
In theory the ambiguity of paleolatitude determination mentioned above can be removed by carrying out the analysis on anomalies of the same age on either side of a particular ridge. An intriguing result of such studies, however, is that in some cases different latitudes of formation are deduced for the same anomaly on either side of the ridge. In that they were formed at the same time and at the same ridge crest, this cannot be so, and the result is giving us yet more information on the geometry or magnetization of the source region. Such a result could be produced by a decay in the intensity of the Earth’s magnetic field or an increase in the number of magnetic excursions or very short-lived reversals as a particular polarity interval progresses, and/or a more complex
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Figure 9 Postulated geometry of normal/reverse magnetization contrasts in the oceanic crust, which could explain the anomalous phase shift of certain linear anomalies.
spreading and volcanic activity, it records both the history of reversals of the Earth’s magnetic field, and the lateral and latitudinal displacement of the crust during the past 160 million years. This record can be played back by measuring the anomalies in the intensity of the Earth’s magnetic field over the oceans at the present day.
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Deep-Sea Drilling Results. Geomagnetic Polarity Timescale. Mid-Ocean Ridge Geochemistry and Petrology. Mid-Ocean Ridge Seismic Structure. Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology. Propagating Rifts and Microplates. Seamounts and Off-Ridge Volcanism. Seismic Structure
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Further Reading
Figure 8 Variation of the magnetic anomaly pattern (A) with geomagnetic latitude (all profiles are N–S; angles refer to magnetic inclination; no vertical exaggeration) and (B) with direction of the profile at fixed latitude (magnetic inclination is 451 in all cases; no vertical exaggeration). (Reproduced from Kearey P and Vine FJ. Global Tectonics. ^ 1996 Blackwell Science, by permission of the publisher.)
Bullard EC and Mason RG (1963) The magnetic field over the oceans. In: Hill MN (ed.) The Sea, vol. 3, pp. 175--217. London: Wiley-Interscience. Harrison CGA (1981) Magnetism of the oceanic crust. In: Emiliani C (ed.) The Sea, vol. 7, pp. 219--239. Wiley: New York. Jones EJW (1999) Marine Geophysics. Chichester: Wiley. Kearey P and Vine FJ (1996) Global Tectonics. Oxford: Blackwell Science. Vacquier V (1972) Geomagnetism in Marine Geology. Amsterdam: Elsevier.
England seamounts of the north–west Atlantic, and the Musician seamounts of the central Pacific. These yield paleomagnetic pole positions that are consistent with other results for the mid-Cretaceous for the North American and Pacific plates, respectively.
Conclusion Thus, because of the dominance of remanent magnetization in the oceanic crust, and the relative simplicity of the way in which it is formed by seafloor
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MANGANESE NODULES D. S. Cronan, Royal School of Mines, London, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1526–1533, & 2001, Elsevier Ltd.
Introduction Manganese nodules, together with micronodules and encrustations, are ferromanganese oxide deposits which contain variable amounts of other elements
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(Table 1). They occur throughout the oceans, although the economically interesting varieties have a much more restricted distribution. Manganese nodules are spherical to oblate in shape and range in size from less than 1 cm in diameter up to 10 cm or more. Most accrete around a nucleus of some sort, usually a volcanic fragment but sometimes biological remains. The deposits were first described in detail in the Challenger Reports. This work was co-authored by J. Murray and A. Renard, who between them initiated
Average abundances of elements in ferromanganese oxide deposits
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the first great manganese nodule controversy. Murray believed the deposits to have been formed by submarine volcanic processes whereas Renard believed that they had precipitated from continental runoff products in sea water. This controversy remained unresolved until it was realized that nodules could obtain their metals from either or both sources. The evidence for this included the finding of abundant nodules in the Baltic Sea where there are no volcanic influences, and the finding of rapidly grown ferromanganese oxide crusts associated with submarine hydrothermal activity of volcanic origin on the Mid-Atlantic Ridge. Subsequently, a third source of metals to the deposits was discovered, diagenetic remobilization from underlying sediments. Thus marine ferromanganese oxides can be represented on a triangular diagram (Figure 1), the corners being occupied by hydrothermal (volcanically derived), hydrogenous (seawater derived) and diagenetic (sediment interstitial water derived) constituents. There appears to be a continuous compositional transition between hydrogenous and diagenetic deposits, all of which are formed relatively slowly at normal deep seafloor temperatures. By contrast, although theoretically possible, no continuous compositional gradation has been reported between hydrogenous and hydrothermal deposits, although mixtures of the two do occur. This may be partly because (1) the growth rates of hydrogenous and hydrothermal deposits are very different with the latter accumulating much more rapidly than the former leading to the incorporation of only limited amounts of the more slowly accumulating hydrogenous material in them, and (2) the temperatures of
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formation of the deposits are different leading to mineralogical differences between them which can affect their chemical composition. Similarly, a continuous compositional gradation between hydrothermal and diagenetic ferromanganese oxide deposits has not been found, although again this is theoretically possible. However, the depositional conditions with which the respective deposits are associated i.e., high temperature hydrothermal activity in mainly sediment-free elevated volcanic areas on the one hand, and low-temperature accumulation of organic rich sediments in basin areas on the other, would preclude much mixing between the two. Possibly they may occur in sedimented active submarine volcanic areas.
Internal Structure The main feature of the internal structure of nodules is concentric banding which is developed to a greater or lesser extent in most of them (Figure 2). The bands represent thin layers of varying reflectivity in polished section, the more highly reflective layers being generally richer in manganese than the more poorly reflective ones. They are thought to possibly represent varying growth conditions.
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Figure 1 Triangular representation of marine ferromanganese oxide deposits.
Figure 2 Concentric banding in a manganese nodule. (Reproduced by kind permission of CNEXO, France.)
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On a microscopic scale, a great variety of structures and textures are apparent in nodules, some of them indicative of postdepositional alteration of nodule interiors. One of the most commonly observed and most easily recognizable is that of collomorphic globular segregations of ferromanganese oxides on a scale of tenths of a millimeter or less, which often persist throughout much of the nodule interior. Often the segregations become linked into polygons or cusps elongated radially in the direction of growth of the nodules. Several workers have also recognized organic structures within manganese nodules. Furthermore, cracks and fissures of various sorts are a common feature of nodule interiors. Fracturing of nodules is a process which can lead to their breakup on the seafloor, in some cases as a result of the activity of benthic organisms, or of bottom currents. Fracturing is an important process in limiting the overall size of nodules growing under any particular set of conditions.
Growth Rates It is possible to assess the rate of growth of nodules either by dating their nuclei, which gives a minimum rate of growth, or by measuring age differences between their different layers. Most radiometric dating techniques indicate a slow growth rate for nodules, from a few to a few tens of millimeters per million years. Existing radiometric and other techniques for nodule dating include uranium series disequilibrium methods utilizing 230Th 231Pa, the 10Be method, the K-Ar method, fission track dating of nodule nuclei, and hydration rind dating. In spite of the overwhelming evidence for slow growth, data have been accumulating from a number of sources which indicate that the growth of nodules may be variable with periods of rapid accumulation being separated by periods of slower, or little or no growth. In general, the most important factor influencing nodule growth rate is likely to be the rate at which elements are supplied to the deposits, diagenetic sources generally supplying elements at a faster rate than hydrogenous sources (Figure 1). Further, the tops, bottoms and sides of nodules do not necessarily accumulate elements at the same rate, leading to the formation of asymmetric nodules in certain circumstances (Figure 3). Differences in the surface morphology between the tops, bottoms and sides of nodules in situ may also be partly related to growth rate differences. The tops receive slowly accumulating elements hydrogenously supplied from seawater and are smooth, whereas the bottoms receive more rapidly accumulating elements diagenetically supplied
Fe, Co from sea water
Sediment surface Mn, Ni, Cu from interstitial waters 1cm Figure 3 Morphological and compositional differences between the top and bottom of a Pacific nodule. (Reproduced with permission from Cronan, 1980.)
from the interstitial waters of the sediments and are rough (Figure 3). The ‘equatorial bulges’ at the sediment–water interface on some nodules have a greater abundance of organisms on them than elsewhere on the nodule surface, suggesting that the bulges may be due to rapid growth promoted by the organisms. It is evident therefore that nodule growth cannot be regarded as being continuous or regular. Nodules may accrete material at different rates at different times and on different surfaces. They may also be completely buried for periods of time during which it is possible that they may grow from interstitial waters at rates different from those while on the surface, or possibly not grow at all for some periods. Some even undergo dissolution, as occurs in the Peru Basin where some nodules get buried in suboxic to reducing sediments.
Distribution of Manganese Nodules The distribution and abundance of manganese nodules is very variable on an oceanwide basis, and can also be highly variable on a scale of a kilometer or less. Nevertheless, there are certain regional regularities in average nodule abundance that permit some broad areas of the oceans to be categorized as containing abundant nodules, and others containing few nodules (Figure 4), although it should always be borne in mind that within these regions local variations in nodule abundance do occur. The distribution of nodules on the seafloor is a function of a variety of factors which include the presence of nucleating agents and/or the nature and age of the substrate, the proximity of sources of elements, sedimentation rates and the influence of organisms. The presence of potential nuclei on the seafloor is of prime importance in determining nodule distribution. As most nodule nuclei are volcanic in origin, patterns of volcanic activity and the
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491
Figure 4 Distribution of mangancese nodules in the oceans (updated from Cronan, 1980 after various sources.) , Areas of nodule coverage; , areas where nodules are locally abundant.
subsequent dispersal of volcanic materials have an important influence on where and in what amounts nodules occur. Other materials can also be important as nodule nuclei. Biogenic debris such as sharks’ teeth, can be locally abundant in areas of slow sedimentation and their distribution will in time influence the abundance of nodules in such areas. As most nuclei are subject to replacement with time, old nodules have sometimes completely replaced their nuclei and have fractured, thus providing abundant nodule fragments to serve as fresh nuclei for ferromanganese oxide deposition. In this way, given sufficient time, areas which initially contained only limited nuclei may become covered with nodules. One of the most important factors affecting nodule abundance on the seafloor is the rate of accumulation of their associated sediments, low sedimentation rates favoring high nodule abundances. Areas of the seafloor where sedimentation is rapid are generally only sparsely covered with nodules. For example, most continental margin areas have sedimentation rates that are too rapid for appreciable nodule development, as do turbiditefloored deep-sea abyssal plains. Low rates of sedimentation can result either from a minimal sediment supply to the seafloor or currents inhibiting its deposition. Large areas in the centers of ocean basins receive minimal sediment input. Under these
conditions substantial accumulation of nodules at the sediment surface is favored. Worldwide Nodule Distribution Patterns
Pacific Ocean As shown in, nodules are abundant in the Pacific Ocean in a broad area, called the Clarion–Clipperton Zone, between about 61N and 201N, extending from approximately 1201W to 1601W. The limits of the area are largely determined by sedimentation rates. Nodules are also locally abundant further west in the Central Pacific Basin. Sediments in the northern part of the areas of abundant nodules in the North Pacific are red clays with accumulation rates of around 1 mm per thousand years whereas in the south they are siliceous oozes with accumulation rates of 3 mm per thousand years, or more. Nodule distribution appears to be more irregular in the South Pacific than in the North Pacific, possibly as a result of the greater topographic and sedimentological diversity of the South Pacific. The nodules are most abundant in basin environments such as those of the south-western Pacific Basin, Peru Basin, Tiki Basin, Penrhyn Basin, and the CircumAntarctic area. Indian Ocean In the Indian Ocean the most extensive areas of nodule coverage are to the south
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of the equator. Few nodules have been recorded in the Arabian Sea or the Bay of Bengal, most probably because of the high rates of terrigenous sediment input in these regions from the south Asian rivers. The equatorial zone is also largely devoid of nodules. High nodule concentrations have been recorded in parts of the Crozet Basin, in the Central Indian Ocean Basin and in the Wharton Basin. Atlantic Ocean Nodule abundance in the Atlantic Ocean appears to be more limited than in the Pacific or Indian Oceans, probably as a result of its relatively high sedimentation rates. Another feature which inhibits nodule abundance in the Atlantic is that much of the seafloor is above the calcium carbonate compensation depth (CCD). The areas of the Atlantic where nodules do occur in appreciable amounts are those where sedimentation is inhibited. The deep water basins on either side of the MidAtlantic Ridge which are below the CCD and which accumulate only limited sediment contain nodules in reasonable abundance, particularly in the western Atlantic. Similarly, there is a widespread occurrence of nodules and encrustations in the Drake Passage–Scotia Sea area probably due to the strong bottom currents under the Circum-Antarctic current inhibiting sediment deposition in this region. Abundant nodule deposits on the Blake Plateau can also be related to high bottom currents. Buried nodules Most workers on the subject agree that the preferential concentration of nodules at the sediment surface is due to the activity of benthic organisms which can slightly move the nodules. Buried nodules have, however, been found in all the oceans of the world. Their abundance is highly variable, but it is possible that it may not be entirely random. Buried nodules recovered in large diameter cores are sometimes concentrated in distinct layers. These layers may represent ancient erosion surfaces or surfaces of nondeposition on which manganese nodules were concentrated in the past. By contrast, in the Peru Basin large asymmetrical nodules get buried when their bottoms get stuck in tenacious suboxic sediment just below the surface layer.
Compositional Variability of Manganese Nodules Manganese nodules exhibit a continuous mixing from diagenetic end members which contain the mineral 10A˚ manganite (todorokite) and are enriched in Mn, Ni and Cu, to hydrogenous end members which contain the mineral d MnO2
(vernadite) and are enriched in Fe and Co. The diagenetic deposits derive their metals at least in part from the recycling through the sediment interstitial waters of elements originally contained in organic phases on their decay and dissolution in the sediments, whereas the hydrogenous deposits receive their metals from normal sea water or diagenetically unenriched interstitial waters. Potentially ore-grade manganese nodules of resource interest fall near the diagenetic end member in composition. These are nodules that are variably enriched in Ni and Cu, up to a maximum of about 3.0% combined. One of the most striking features shown by chemical data on nodules are enrichments of many elements over and above their normal crustal abundances (Table 1). Some elements such as Mn, Co, Mo and Tl are concentrated about 100-fold or more; Ni, Ag, Ir and Pb are concentrated from about 50- to 100-fold, B, Cu, Zn, Cd, Yb, W and Bi from about 10 to 50-fold and P, V, Fe, Sr, Y, Zr, Ba, La and Hg up to about 10-fold above crustal abundances.
Regional Compositional Variability
Pacific Ocean In the Pacific, potentially ore-grade nodules are generally confined to two zones running roughly east–west in the tropical regions, which are well separated in the eastern Pacific but which converge at about 1701–1801W (Figure 5). They follow the isolines of intermediate biological productivity, strongly suggestive of a biological control on their distribution. Within these zones, the nodules preferentially occupy basin areas near or below the CCD. Thus they are found in the Peru Basin, Tiki Basin, Penrhyn Basin, Nova Canton Trough area, Central Pacific Basin and Clarion– Clipperton Zone (Figure 5). Nodules in all these areas have features in common and are thought to have attained their distinctive composition by similar processes. The potentially ore-grade manganese nodule field in the Peru Basin, centered at about 71–81S and 901W (Figure 5), is situated under the southern flank of the equatorial zone of high biological productivity on a seafloor composed of pelagic brown mud with variable amounts of siliceous and calcareous remains. Nodules from near the CCD at around 4250 m are characterized by diagenetic growth and are enriched in Mn, Ni and Cu, whereas those from shallower depth are characterized mainly by hydrogenous growth. The Mn/Fe ratio increases from south to north as productivity increases, whereas the Ni and Cu contents reach maximum values in the middle of the area where Mn/Fe ratios are about 5.
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50 Figure 5 Approximate limits of areas of nickel- and copper-rich nodules in the subequatorial Pacific referred to in the text (productivity isolines if g C m2 y1).
In the Tiki Basin there is also an increase in the Mn/Fe ratio of the nodules from south to north. All Ni þ Cu values are above the lower limit expected in diagenetically supplied material. The Penrhyn Basin nodules fall compositionally within the lower and middle parts of the Mn/Fe range for Pacific nodules as a whole. However, nodules from the northern part of the Basin have the highest Mn/Fe ratios and highest Mn, Ni and Cu concentrations reflecting diagenetic supply of metals to them, although Ni and Cu decrease slightly as the equator is approached. Superimposed on this trend are variations in nodule composition with their distance above or below the CCD. In the Mn-, Ni-, and Cu-rich nodule area, maximum values of these metals in nodules occur within about 200 m above and below the CCD. The latititudinal variation in Mn, Ni and Cu in Penrhyn Basin nodules may be due to there being a hydrogenous source of these metals throughout the Basin, superimposed on which is a diagenetic source of them between about 21 and 61S at depths near the CCD, but less so in the very north of the Basin (0–21S) where siliceous sedimentation prevails under highest productivity waters. In the Nova Canton Trough area, manganese concentrations in the nodules are at a maximum between the equator and 2.51S, where the Mn/Fe ratio is also highest. Manganese shows a tendency to
decrease towards the south. Nickel and copper show similar trends to Mn, with maximum values of these elements being centered just south of the equator at depths of 5300–5500 m, just below the CCD. In the central part of the Central Pacific Basin, between the Magellan Trough and the Nova Canton Trough, diagenetic nodules are found associated with siliceous ooze and clay sedimentation below the CCD. Their Ni and Cu contents increase southeastwards reaching a maximum at about 2.51–31N and then decrease again towards the equator where productivity is highest. The Clarion–Clipperton Zone deposits rest largely on slowly accumulated siliceous ooze and pelagic clay below the CCD. The axis of highest average Mn/ Fe ratio and Mn, Ni and Cu concentrations runs roughly southwest–northeast with values of these elements decreasing both to the north and south as productivity declines respectively to the north and increases towards the equatorial maximum in the south. Indian Ocean In the Indian Ocean, Mn-, Ni-, and Cu-rich nodules are present in the Central Indian Ocean Basin between about 51 and 151S. They are largely diagenetic in origin and rest on siliceous sediments below the CCD under high productivity waters. The deposits show north–south compositional
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variability with the highest grades occurring in the north.
Atlantic Ocean In the Atlantic Ocean, diagenetic Mn-, Ni-, and Cu-rich nodules occur most notably in the Angola Basin and to a lesser extent in the Cape/Agulhas Basin and the East Georgia Basin. These three areas have in common elevated biological productivity and elevated organic carbon contents in their sediments, which coupled with their depth near or below the CCD would help to explain the composition of their nodules. However, Ni and Cu contents are lower in them than in areas of diagenetic nodules in the Pacific and Indian Oceans.
Economic Potential Interest in manganese nodules commenced around the mid-1960s and developed during the 1970s, at the same time as the Third United Nations Law of the Sea Conference. However, the outcome of that Conference, in 1982, was widely regarded as unfavorable for the mining industry. This, coupled with a general downturn in metal prices, resulted in a lessening of mining company interest in nodules. About this time, however, several governmentbacked consortia became interested in them and this work expanded as evaluation of the deposits by mining companies declined. Part 11 of the 1982 Law of the Sea Convention, that part dealing with deepsea mining, was substantially amended in an agreement on 28 July 1994 which ameliorated some of the provisions relating to deep-sea mining. The Convention entered into force in November 1994. During the 1980s interest in manganese nodules in exclusive economic zones (EEZs) started to increase. An important result of the Third Law of the Sea Conference, was the acceptance of a 200-nauticalmile EEZ in which the adjacent coastal state could claim any mineral deposits as their own. The nodules found in EEZs are similar to those found in adjacent parts of the International Seabed Area, and are of greatest economic potential in the EEZs of the South Pacific. At the beginning of the twenty-first century, the out-look for manganese nodule mining remains rather unclear. It is likely to commence some time in this century, although it is not possible to give a precise estimate as to when. The year 2015 has been suggested as the earliest possible date for nodule mining outside of the EEZs. It is possible, however, that EEZ mining for nodules might commence earlier
if conditions were favorable. It would depend upon many factors; economic, technological, and political.
Discussion A model to explain the compositional variability of nodules in the Penrhyn Basin can be summarized as follows. Under the flanks of the high productivity area, reduced sedimentation rates near the CCD due to calcium carbonate dissolution enhance the content of metal-bearing organic carbon rich phases (fecal material, marine snow, etc.) in the sediments, the decay of which drives the diagenetic reactions that in turn promote the enrichment of Mn, Ni, and Cu in the nodules via the sediment interstitial waters. Away from the CCD, organic carbon concentrating processes are less effective. Further south as productivity declines, there is probably insufficient organic carbon supplied to the seafloor to promote the formation of diagenetic nodules at any depth. Under the equator, siliceous ooze replaces pelagic clay as the main sediment builder at and below the CCD, and when its rate of accumulation is high it dilutes the concentrations of organic carbon-bearing material at all depths to levels below that at which diagenetic Mn, Ni, and Cu rich nodules can form. To a greater or lesser extent, this model can account for much of the variability in nodule composition found in the other South Pacific areas described, although local factors may also apply. In the Peru Basin, as in the Penrhyn Basin, diagenetic Mn-, Ni-, and Cu-rich nodules are concentrated near the CCD and their Ni and Cu contents reach a maximum south of the highest productivity waters. In the Tiki Basin, the greatest diagenetic influences are also found in the north of the Basin. As the South Pacific basins deepen to the west, the areas of diagenetic nodules tend to occur below the CCD as, for example, in the Nova Canton Trough area. This may be because the settling rates of large organic particles are quite fast in the deep ocean. Probably only limited decay of this material takes place between it settling through the CCD and reaching the seafloor, and enough probably gets sedimented to extend the depth of diagenetic nodule formation to well below the CCD under high productivity waters where there is limited siliceous sediment accumulation. In the North Pacific, the trends in nodule composition in relation to the equatorial zone are the mirror image of those in the south. Thus in both the Central Pacific Basin and the Clarion–Clipperton Zone the highest nodule grades occur in diagenetic nodules on the northern flanks of the high
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MANGANESE NODULES
productivity area and decline both to the north and south. The general model erected to explain the Penrhyn Basin nodule variability thus probably applies, at least in part, to these areas also. The model also has some applicability in the Indian Ocean but less in the Atlantic. In the Indian Ocean, diagenetic nodules associated with sediments containing moderate amounts of organic carbon occur resting on siliceous ooze to the south of the equatorial zone in the Central Indian Ocean Basin. Farther to the south these nodules give way to hydrogenous varieties resting on pelagic clay. However, in the north the changes in nodule composition that might be expected under higher productivity waters do not occur, probably because terrigenous sedimentation becomes important in those areas which in turn reduces the Mn, Ni, and Cu content of nodules. In the Atlantic, the influence of equatorial high productivity on nodule composition that is evident in the Pacific is not seen, mainly because the seafloor in the equatorial area is largely above the CCD. Where diagenetic nodules do occur, as in the Angola, Cape and East Georgia Basins, productivity is also elevated, but the seafloor is near or below the CCD leading to reduced sedimentation rates.
Conclusions Manganese nodules, although not being mined today, are a considerable resource for the future. They consist of ferromanganese oxides variably enriched in Ni, Cu, and other metals. They generally accumulate around a nucleus and exhibit internal layering on both a macro- and microscale. Growth rates are generally slow. The most potentially economic varieties of the deposits occur in the subequatorial Pacific under the flanks of the equatorial
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zone of high biological productivity, at depths near the CCD. Similar nodules occur in the Indian Ocean under similar conditions.
See also Authigenic Deposits. Hydrothermal Vent Biota. Hydrothermal Vent Ecology. Hydrothermal Vent Fluids, Chemistry of.
Further Reading Cronan DS (1980) Underwater Minerals. London: Academic Press. Cronan DS (1992) Marine Minerals in Exclusive Economic Zones. London: Chapman and Hall. Cronan DS (ed.) (2000) Handbook of Marine Mineral Deposits. Boca Raton: CRC Press. Cronan DS (2000) Origin of manganese nodule ‘ore provinces’. Proceedings of the 31st International Geological Congress, Rio de Janero, Brazil, August 2000. Earney FC (1990) Marine Mineral Resources. London: Routledge. Glasby GP (ed.) (1977) Marine Manganese Deposits. Amsterdam: Elsevier. Halbach P, Friedrich G, and von Stackelberg U (eds.) (1988) The Manganese Nodule Belt of the Pacific Ocean. Stuttgart: Enke. Nicholson K. Hein J, Buhn B, Dasgupta S (eds.) (1997) Manganese Mineralisation: Geochemistry and Mineralogy of Terrestrial and Marine Deposits. Geological Society Special Publication 119, London. Roy S (1981) Manganese Deposits. London: Academic Press. Teleki PG, Dobson MR, Moore JR, and von Stackelberg U (eds.) (1987) Marine Minerals: Advances in Research and Resource Assessment. Dordrecht: D. Riedel.
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MANGROVES M. D. Spalding, UNEP World Conservation Monitoring Centre and Cambridge Coastal Research Unit, Cambridge, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 3, pp 1533–1542, & 2001, Elsevier Ltd.
highly varied salinities, often approaching hypersaline conditions. Soils may be shallow, but even where they are deep they are usually anaerobic within a few millimeters of the soil surface. Many mangrove species show one or more of a range of physiological, morphological or life-history adaptations in order to cope with these conditions.
Definition
Coping with Salt
The term mangrove is used to define both a group of plants and also a community or habitat type in the coastal zone. Mangrove plants live in or adjacent to the intertidal zone. Mangrove communities are those in which these plants predominate. Other terms for these communities include coastal woodland, intertidal forest, tidal forest, mangrove forest, mangrove swamp and mangal. The word mangrove can be clearly traced to the Portugese word ‘mangue’ and the Spanish word ‘mangle’, both of which are actually used in the description of the habitats, rather than the plants themselves, but still have been joined to the English word ‘grove’ to give the word ‘mangrove.’ It has been suggested that the original Portugese word has been adapted from a similar word used locally by the people of Senegal, however an alternative derivation may be the word ‘manggimanggi’, which is still used in parts of eastern Indonesia to describe one genus (Avicennia).
All mangroves are able to exclude most of the salt in sea water from their xylem. The exact mechanisms for this remain unclear, but it would appear to be an ultrafiltration process operating at the endodermis of the roots. One group, which includes Bruguiera, Lumnitzera, Rhizophora and Sonneratia, is highly efficient in this initial salt exclusion and shows only minor further mechanisms for salt secretion. A second group, which includes Aegialitis, Aegiceras and Avicennia appear to be less efficient at this initial salt exclusion and hence also need to actively secrete salt from their leaves. This is done metabolically, using special salt glands on the leaf surface. The salt evaporates, leaving crystals which may be washed or blown off the leaf surface. In these latter species such exuded salt is often visible on the leaf surface (Figure 1).
Mangrove Species Mangrove plants are not a simple taxonomic group, but are largely defined by the ecological niche where they live. The simplest definition describes a shrub or tree which normally grows in or adjacent to the intertidal zone and which has developed special adaptations in order to survive in this environment. Using such a definition a broad range of species can be identified, coming from a number of different families. Although there is no consensus as to which species are, or are not, true mangroves, there is a core group of some 30–40 species which are agreed by most authors. Furthermore, these ‘core’ species are the most important, both numerically and structurally, in almost all mangrove communities. Table 1 lists a large range of mangrove species (of tree, shrub, fern, and palm), and highlights those which might be regarded as core species. All of these plants have adapted to a harsh environment, with regular inundation of the soil and
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Anaerobic Soils
The morphological feature for which mangroves are best known is the development of aerial roots. These have developed in most mangrove species in order to cope with the need for atmospheric oxygen at the absorbing surfaces and the impossibility of obtaining such oxygen in an anaerobic and regularly inundated environment. Various types of roots are illustrated in Figure 1. The stilt root, exemplified by Rhizophora (Figure 1B) consists of long branching structures which arch out away from the tree and may loop down to the soil and up again. Such stilt roots also occur in Bruguiera and Ceriops although in older specimens they fuse to the trunk as buttresses. They also occur sporadically in other species, including Avicennia. A number of unrelated groups have developed structures known as pneumatophores which are simple upward extensions from the horizontal root into the air above. These are best developed in Avicennia and Sonneratia (Figure 1C), the former typically having narrow, pencil-like pneumatophores, the
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Table 1
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List of mangrove species: species in bold typeface are those which are considered ‘core’ species
Family
Species
Pteridaceae
Acrostichum aureum Acrostichum danaeifolium Acrostichum speciosum Aegialitis annulata Aegialitis rotundifolia Pelliciera rhizophorae Camptostemon philippensis Camptostemon schultzii Heritiera fomes Heritiera globosa Heritiera littoralis Diospyros ferrea Aegiceras corniculatum Aegiceras floridum Cynometra iripa Mora oleifera Conocarpus erectus Laguncularia racemosa Lumnitzera littorea Lumnitzera racemosa Lumnitzera rosea Pemphis acidula Osbornia octodonta Sonneratia alba Sonneratia apetala Sonneratia caseolaris Sonneratia griffithii Sonneratia lanceolata Sonneratia ovata Sonneratia gulngai Sonneratia urama Bruguiera cylindrica Bruguiera exaristata Bruguiera gymnorrhiza Bruguiera hainesii
Plumbaginaceae Pellicieraceae Bombacaceae Sterculiaceae
Ebenaceae Myrsinaceae Caesalpiniaceae Combretaceae
Lythraceae Myrtaceae Sonneratiaceae
Rhizophoraceae
Family
latter with secondary thickening so that they can become quite tall and conical. One adaptation on the theme of pneumatophores is that of root knees where more rounded knobs are observed to extend upwards from the roots. In Xylocarpus mekongensis these are simply the result of localized secondary cambial growth, but in Bruguiera (Figure 1D) and Ceriops they are the result of a primary looping growth. In these species branching may also occur on these root knees. Buttress roots are a common adaptation of many tropical trees, but in Xylocarpus granatum (Figure 1E) and to some degree in Heritiera such flange-like extensions of the trunk continue into plank roots which are vertically extended roots with a sinuous plank-like form extending above the soil. The surfaces of the aerial roots are amply covered with porous lenticels to enable gaseous exchange, and the internal structure of the roots is highly
Euphorbiaceae Meliaceae
Avicenniaceae
Acanthaceae Bignoniaceae Rubiaceae Arecaceae
Species Bruguiera parviflora Bruguiera sexangula Ceriops australis Ceriops decandra Ceriops tagal Kandelia candel Rhizophora apiculata Rhizophora harrisonii Rhizophora mangle Rhizophora mucronata Rhizophora racemosa Rhizophora samoensis Rhizophora stylosa Rhizophora lamarckii Rhizophora selala Excoecaria agallocha Excoecaria indica Aglaia cucullata Xylocarpus granatum Xylocarpus mekongensis Avicennia alba Avicennia bicolor Avicennia germinans Avicennia integra Avicennia marina Avicennia officinalis Avicennia rumphiana Avicennia schaueriana Acanthus ebracteatus Acanthus ilicifolius Dolichandrone spathacea Tabebuia palustris Scyphiphora hydrophyllacea Nypa fruticans
adapted, with large internal gas spaces, making up around 40% of the total root volume in some species. It is further widely accepted that there must be some form of ventilatory mechanism to aid gaseous exchange. A system of tidal suction is the probable mechanism in most species: during high tides, oxygen is used by the plant, while carbon dioxide is readily absorbed in the sea water, leading to reduced pressure within the roots. As the tide recedes and the lenticels open, water is then sucked into the roots. Seeds and Seedlings
Establishment of new mangrove plants in the unstable substrates and regular tidal washing of the mangrove environment presents a particular evolutionary challenge. All mangroves are dispersed by water and particular structures in the seed or the fruit are adapted to support flotation. In a number of
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groups a degree of vivipary is observed which is unusual in most nonmangroves. The Rhizophoraceae have developed this to its fullest extent and here the embryo grows out of the seed coat and then out of the fruit while still attached to the parent plant, so that the propagule which is eventually released is actually a seedling rather than a seed (Figure 1F). In a number of other groups, including Aegiceras, Avicennia, Nypa and Pelliciera cryptovivipary exists in which the embryo emerges from the seed coat, but not the fruit, prior to abcission. Longevity of seedlings is clearly important for many species. Most species are able to survive (float and remain viable) for over a month, whereas some Avicennia propagules have been shown to remain viable for over a year while in salt water.
Distribution and Biogeography As a result of their restriction to intertidal areas, mangroves are limited in global extent (Figure 2), and are, in fact, one of the most globally restricted of all forest types. Figure 2 clearly shows the absolute limits to mangrove distribution. Mangroves are largely confined to the regions between 301 north and south of the equator, with notable extensions beyond this to the north in Bermuda (321200 N) and Japan (311220 N), and to the south in Australia (381450 S), New Zealand (381030 S) and South Africa (321590 S). Within these confines they are widely distributed, although their latitudinal development is restricted along the western coasts of the Americas and Africa. In the Pacific Ocean natural mangrove
(A)
(C)
(B)
(D)
Figure 1 Mangrove adaptations: (A) salt crystals secreted onto the surface of a leaf, Avicennia; (B) stilt roots of Rhizophora; (C) pneumatophores in Sonneratia; (D) root knees in Bruguiera; (E) plank roots in Xylocarpus; (F) Rhizophora propagule.
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1. There is a distinct region of very high mangrove diversity, sometimes referred to as the ‘diversity anomaly’, centered over south-east Asia. 2. Away from this high diversity region mangroves generally show relatively even levels of low diversity, although there is smaller peak of diversity around southern Central America. 3. There is a wide area of the central and western Pacific Ocean from 120 to 1601W where mangroves do not occur. 4. Even in the area of highest mangrove diversity there is a very rapid latitudinal decline in species numbers away from the tropics. One further observation, which is not fully illustrated in the figures, concerns the division of the global mangrove flora into two highly distinct subregions. An eastern group (sometimes known as the Indo-West Pacific) forms one vast and contiguous block stretching from the Red Sea and East Africa to the central Pacific. This group has a totally different species composition from the western group (the Atlantic–East Pacific or Atlantic–Caribbean–East Pacific), which includes both Pacific and Atlantic shores of the Americas, the Caribbean and the shores of West Africa. A number of these patterns are explored more fully, below. Latitudinal Patterns
Mangroves limits are closely correlated to minimum temperature requirements. There is only one genus (Avicennia) which survives in environments where frosts may occur, but many species appear to have their latitudinal limits set by less extreme cold temperatures; air temperatures of 51C appear inimical to most mangrove species. Sea-surface temperatures may be more important than air temperatures for some species. The 241C mean annual isotherm appears to be the minimum water temperature tolerated by mangroves in most areas, although this
Figure 1 Continued
communities are limited to western areas, and they are absent from many Pacific islands. In all, an estimated 114 countries and territories have mangroves, however for many nations the total area is very small indeed, and the total global area of these forests is only 181 000 km2. Table 2 provides a summary of total mangrove areas by region. Although these statistics suggest a relatively wide distribution, the distribution of individual species within these areas is clearly far more restricted, and Figure 3 provides a plot of mangrove biodiversity patterns. A number of points of particular interest are clearly illustrated.
Table 2
Total mangrove area by region
Region
South and south-east Asia Australasia The Americas West Africa East Africa and the Middle East Total area
Area (km2)
75 18 49 27 10
173 789 096 995 024
Proportion of global total 41.5% 10.4% 27.1% 15.5% 5.5%
181 077
Data calculated from best available national sources in Spalding et al. (1997).
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Figure 3 A global map of mangrove diversity plotting contours of equal diversity (1–5 species, 6–10, 11–15, 16–20, 21–25, 26–30, 31–35, 36–40, and 41–45). (Reproduced with permission from Spalding (1998).)
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Figure 2 The global distribution of mangrove forests (data kindly provided by the UNEP World Conservation Monitoring Centre).
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minimum is closer to 271C on the north Atlantic coasts of America and Africa, and may be much lower in areas such as southern Japan. The relatively low-latitude limits to mangrove in Peru and Angola are probably related to the cold water currents which affect these coastlines. Eastern and Western Floras
The division of mangroves into two distinct floras is almost complete at the species level. Of the species listed i
