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 4: N - R 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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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.
(c) 2011 Elsevier Inc. All Rights Reserved.
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.
(c) 2011 Elsevier Inc. All Rights Reserved.
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
(c) 2011 Elsevier Inc. All Rights Reserved.
513
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
(c) 2011 Elsevier Inc. All Rights Reserved.
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.
NEKTON W. G. Pearcy, Oregon State University, Corvallis, OR, USA R. D. Brodeur, Northwest Fisheries Science Center, Newport, OR, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Marine nekton are the swimmers in the sea, as opposed to plankton that drift with currents. Although the main criterion is swimming ability, most definitions of nekton, as with plankton, are not precise. They are based on swimming speeds, lengths, Reynolds number, shape to reduce drag, or simply those animals too swift or agile to be captured in plankton nets. Nekton include fishes, cephalopods, seabirds, marine mammals, reptiles, and crustaceans. They inhabit all the ecological zones of the ocean, from the epipelagic near-surface waters to the deep-sea abyss, including both pelagic and epibenthic animals. Methods of assessing the distribution and abundances of nekton include nets, baited hooks or traps, acoustics, visual systems, lasers, egg/larval surveys, tagging/recapture, and food habits of their predators. Because this edition includes articles on the biology of fishes, fisheries, birds, and marine mammals, the emphasis here will be on the smaller nekton, or micronekton (usually 2–20 cm in length). Although most micronekton are not harvested, krill (euphausiids) and lanternfishes are sometimes fished commercially. Some problems for nektonic animals in the sea include overcoming the forces of drag while moving through a viscous medium, counteracting sinking through buoyancy adaptations, feeding, and avoiding being eaten. The reader is referred to articles on biology of fishes that address some of these subjects. The major taxonomic groups of micronekton are fishes, cephalopods, and large crustaceans. They are important components of marine ecosystems and vital links between zooplankton and higher trophic levels. Many are carnivorous; some are herbivores. As a group they are important to us as food, either directly through fisheries or indirectly as food for commercially important species. Fishes, squids, and euphausiids are important prey for seabirds, larger fishes, and marine mammals. Oceanic and coastal micronekton are important, but poorly understood,
intermediate links in the food webs from zooplankton to apex predators for many marine ecosystems including the tropical Pacific, the Gulf of Alaska, and the Barents Sea (Figure 1). Euphausiids are especially important as prey for many species of fishes, seabirds, and marine mammals. In the Antarctic, krill (mainly Euphausia superba; Figure 2) are a major food for whales, penguins, and some pinnipeds. Although most euphausiids are herbivorous, most of the micronektonic fishes and squids are carnivorous, preying on euphausiids, copepods, amphipods, and other zooplankton. Distributions
Many studies have shown that the distributions of mesopelagic micronekton, like zooplankton, generally coincide with specific water masses in the ocean, often forming faunal assemblages. However, some species are found in all oceans while others have distributions restricted to waters of the continental shelf or slope. Species richness of the micronektonic fauna usually increases from high to low latitudes. The biology and morphology of micronekton differ among the different pelagic realms: epipelagic (0–200 m), mesopelagic (200–1000 m), bathy/abyssopelagic (41000 m), and epibenthic. Epipelagic micronekton include the juvenile phases of many species, as well as the mature stages of small fishes such as Engraulidae (anchovies), Clupeidae (herrings, sardines), and Osmeridae (smelts) (see Figure 1). These epipelagic species are most common over productive regions of the continental shelves and coastal upwelling regions where they are often subjected to commercial fisheries.
Midwater Micronekton Many families of fishes, squids, and crustaceans are considered micronekton (Figure 3). Myctophidae (lanternfishes) and Gonostomatidae (bristlemouths) are common families of mesopelagic fishes. It has been estimated that the worldwide biomass of mesopelagic fishes alone is over 900 million metric tons. Many of these animals have distributions in oceanic (waters beyond the continental shelves) in all the world’s oceans. Generally, mesopelagic micronekton are most abundant over the continental slopes and their numbers decrease both inshore and offshore. Diel migrations also provide an abundance of food resource for shallow-water or coastal fishes.
(c) 2011 Elsevier Inc. All Rights Reserved.
1
2
NEKTON
Tropical Pacific Primary producers
Zooplankton
Micronekton
Macronekton
Apex predators
Sea turtles Meroplankton
Epipelagic
Seabirds Marlin and sailfish
Euphausids Amphipods
Mesopelagic migrant
Copepods
Mesopelagic nonmigrant
Skipjack
Dolphins
Phytoplankton
DOM
Chaetognath Salps
Picoplankton Detritus
Yellow fin
Young tuna and scombrids
Bathypelagic migrant
Piscivorous fish
Bathypelagic nonmigrant
Large squids
Blue shark Other sharks Bigeye
Microzooplankton Larvaceans
Albacore
Swordfish Killer whales
Gulf of Alaska Euphausiids
Pollock Eulachon Pacificcod
Copepods Capelin Chaetognath Phytoplankton
Other
Arrowtooth flounder Halibuth
Pteropods
Purpoises Toothed whales Baleen whales Seabirds
Rockfish Juvenile gadids Larvaceans
Steller sea lion Other
Shrimps
Harbor seals
Flathead sole
Other
Barents Sea
Euphausiids
Herring
Cod
Seabirds
Capelin Copepods Phytoplankton
Polar cod
Other gadids
Baleen whales
Amphipods Juv. gadids Pinnipeds Other
Shrimps
Fatfish
Figure 1 Simplified food web for tropical Pacific, Gulf of Alaska, and Barents Sea showing the intermediate position of micronekton in the food webs. Adapted from Ciannelli L, Hjermann DO, Lehodey P, Ottersen G, Duffy-Anderson JT, and Stenseth NC (2005) Climate forcing, food web structure, and community dynamics in pelagic marine ecosystems. In: Belgrano A, Scharler UM, Dunne J, and Ulanowicz RE (eds.) Aquatic Food Webs, an Ecosystem Approach, pp. 143–169. Oxford, UK: Oxford University Press.
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NEKTON
Figure 2 Euphausiid swarm. Photo courtesy of Bruce Robison, Monterey Bay Aquarium Research Institute (MBARI).
Off Oregon, rockfishes (Sebastes spp.) often prey on euphausiids and sergestid shrimp and myctophid fishes that are carried by currents inshore and then trapped against the shallow seafloor or offshore banks or seamounts during daytime. Off Hawaii, the animals forming layers are impinged along the sides or on the tops of seamounts where they are easy prey (Figure 4). In Hawaiian coastal waters, a mesopelagic boundary layer community consists of a distinct resident community of micronekton distributed along a narrow band where the upper slope meets the oceanic realm. Based on acoustical measurements from moorings, it exhibits both diel vertical and horizontal migrations. Remarkable inshore migrations at night extend within 1 km from shore. Again, this provides a trophic link between neritic and oceanic systems. Diel Vertical Migrations
Many species of epipelagic and mesopelagic micronekton undertake diel, nocturnal, or daily vertical migrations, swimming toward the surface at night and descending into deeper water during periods of daylight. This common and unique behavior of mesopelagic micronekton is shown in Figure 4, where a distinct sound-scattering layer migrates toward the surface during twilight, while other layers do not migrate. Interestingly, not all individuals migrate even within the same species. Some are basically nonmigratory. Some studies have observed diel migrations of sound scatterers between the depths of 500 and 1000 m. However, there is little evidence for migrations of animals living below about 1000 m in the open ocean at bathy/abyssopelagic depths. These diel migrations were first observed in the open ocean by physicists who were perplexed by
3
sound-scattering layers at mid-depths during the day that looked like false bottoms, and often moved toward the surface at night. They called them ‘deepscattering layers’. Different animals reflect sound depending on the frequency of sound used and the sound velocity and density contrast of the animals. We know that the animals that reflect the sound (10–50 kHz) in these layers are usually fishes or siphonophores, often with gas-filled swimbladders or floats that effectively resonate sound. In the open ocean, migrations are generally from mesopelagic depths (200–1000 m) into epipelagic waters, or even to the surface, at night. Because of the huge biomass involved, these migrations of micronekton account for the largest daily movement of biomass on Earth (except for the masses of Homo sapiens that commute daily!). Sunlight is the primary sensory cue that initiates or guides diel vertical migrations. This conclusion is supported by the conformity of the vertical movements of scattering layers and specific isolumes or light intensities, as well as responses to artificial light or solar eclipses. However, endogenous rhythms or changes of light intensity have also been suggested as triggers to initiate migrations. Adaptive significance of diel vertical migrations Why migrate vertically? The main theories concern the predator following prey and the latter avoiding predation. Food is available in greater quantity in shallower waters that are rich with zooplankton but these sunlit waters are dangerous during the daytime because of large visual predators. Hence this behavior is thought to be an energetic trade-off between the risk of predation and feeding. Vulnerability to large visual predators is reduced at night in near-surface water and in deep, dimly lit water during the daytime. Other theories involve an energy bonus for animals feeding in warmer nearsurface waters and digesting and assimilating food in deeper cooler waters. Mid-water fishes often migrate vertically in the Antarctic even when the water column has near-uniform temperature and their major prey, E. superba, remains near the surface and does not migrate. This suggests that avoidance of visual predators is a major selection factor for diel vertical migrations in these waters. These ubiquitous vertical migrations in the open ocean contribute to the vertical transport of energy and elements. Carbon in zooplankton, as well as anthropogenic materials such as pesticides or radioisotopes, are ingested by micronekton in near-surface waters and are defecated in deeper water. This rapid transport of materials has been called a ‘biological pump’.
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4
NEKTON
Figure 3 Photos of in situ mesopelagic fishes, squids, and shrimp taken from a remotely operated vehicle (ROV) Tiburon. All images & MBARI. From top, left to right: lanternfish Stenobrachius leucopsarus, bristlemouth Cyclothone signata, viperfish Chauliodus macouni, squids Galiteuthis, Gonatidae, and shrimp Sergestes similis. All images & MBARI.
Bioluminescence
Many species of oceanic micronekton, especially in the meso/bathypelagic regions, are bioluminescent. These include most families of euphausiids, many pelagic mysids, shrimps, fishes, and squids. Bioluminescence is produced by simple or complex light
organs, photophores, or luminous secretions, either by symbiotic bacteria or intrinsic organs under nervous control. The color of most bioluminescence in mesopelagic animals is blue-green, c. 470–490 nm. Notable exceptions are some stomiatoid fishes that emit a red light, a frequency that is invisible to most
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NEKTON
Depth (m)
Migratory layer ≈ 120 m
9 Oct. 2004
0
5
250
500
Daytime layer ≈ 550 m
750
18:10
18:37
19:04 Time (local)
19:31
Nonmigratory layer ≈ 550 m
Figure 4 A 38-kHz echogram off Hawaii showing the complexity of sound-scattering layers, with a layer migrating toward the surface at twilight and nonmigratory layers near the surface and along the flanks of the seamount.
deep-water nekton – it functions as an underwater snooperscope. Many of the mesopelagic fishes, squids, and crustaceans have photophores arrayed along the ventral portions of their bodies (see fishes in Figure 3). This is thought to provide a ventral ‘countershading’ or counter-illumination that reduces the silhouette of their bodies from predators that lurk below. This theory is supported by experiments that have shown that the intensity of light emitted by the photophores closely matches downwelling light from above. Other bioluminescent organs are thought to be lures to attract prey or to promote schooling or mate recognition. Bioluminescence occurs in the ocean, even down to the deepest depths. If a light meter is lowered to measure light intensity, downwelling light from the sun decreases exponentially, but then light intensity may increase with depth at mid-depths from flashes of bioluminescent organisms. Because bioluminescence is the predominant source of light below 500–1000 m, fishes and crustaceans in the deep ocean usually have either jet-black (fishes) or crimson (crustaceans) pigments, presumably to absorb bioluminescent flashes. Here silvery or reflective bodies, as found in most epipelagic fishes, would reflect these flashes and stand out as mirrors that would attract predators. Life History and Biology
In the large volume that comprises the mid-water domain of the world’s oceans, finding and recognizing mates to reproduce with can be difficult, especially when overall abundances are low. Many fishes depend on random encounters of males and females, perhaps recognizing their own kind by their
photophore patterns. Some male anglerfishes have evolved a bizarre behavior where they actually attach as a dwarf form to a female once they locate them. They then draw nutrition from the female and continue in this parasitic relationship for the rest of their lives. Reproductive patterns are not well known in many mid-water nekton. Myctophids are known to spawn year-round in the tropics but spawn during a single season at higher latitudes. Many myctophids undertake extended seasonal north–south migrations between feeding and spawning grounds, as seen in epipelagic fishes. Most larvae and early juveniles occur in the epipelagic zone, are often nonmigratory, and are found progressively deeper as they grow older, metamorphosing into the adult stage in the mesopelagic zone. They usually mature after 1–3 years but their typical life spans are not generally known. With the exception of Antarctic krill, which can live to be about 7 years old, most nektonic crustaceans and squid probably live only a few years.
Food Habits Feeding modes in midwater nekton can range from filter-feeding to highly carnivorous. Most deep-water animals are carnivorous (note the viperfish in Figure 5). Numerous morphological or behavioral adaptations have evolved to overcome the paucity of food in the deep ocean. Bioluminescence is used to both locate prey, as discussed previously, and attract prey. Anglerfishes often use lighted ‘lures’ suspended over the mouth by fin rays modified to work as ‘fishing poles’ to attract gullible prey. Most species have upward-facing eyes and mouths to locate overhead prey that may be backlit against the
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6
NEKTON
Figure 5 Head-on view of the viperfish, Chauliodus macouni. Image & 2006 MBARI.
Figure 6 Whiptail gulper, Saccopharynx. Note large mouth and extended stomach.
downwelling light. In many cases, the jaws are unhinged or extremely large relative to the size of the organism to swallow large prey. In an extreme example, the whiptail gulper fish (Figure 6) possesses a mouth and stomach that are highly extensible to consume prey bigger than themselves similar to snakes, and then gradually digest the prey item until the next meal comes along, perhaps months later. Given the large number of piscivorous predators in these regions, it is not surprising that many micronekton have developed adaptations for predator avoidance. Their bioluminescent countershading and light absorptive pigmentation reduce their visibility to most predators. Bioluminescent flashes may also confuse predators. Some animals align themselves vertically in the water column to minimize their
cross-sectional area visible to predators from below. The squid Galiteuthis (see Figure 3), which has a transparent body, often holds its opaque arms and tentacles directly over the large downward-oriented light organs on its head, presumably to minimize the silhouette of its arms and tentacles by counterillumination. Other elongate fishes curl up into a ball to mimic other unpalatable species or appear larger than they actually are to a predator. Squids have developed a different strategy in that they can release a black or luminous ink that may temporarily confuse their predators or act as ‘decoys’, allowing their getaway from predators.
See also Cephalopods.Crustacean Fisheries
Further Reading Benoit-Bird KJ and Au WWL (2006) Extreme diel horizontal migrations by a tropical nearshore resident micronekton community. Marine Ecology Progress Series 319: 1--14. Brodeur RD and Yamamura O (2005) Micronekton of the North Pacific. PICES Scientific Report 30: 115pp. Ciannelli L, Hjermann DO, Lehodey P, Ottersen G, DuffyAnderson JT, and Stenseth NC (2005) Climate forcing, food web structure, and community dynamics in pelagic
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NEKTON
marine ecosystems. In: Belgrano A, Scharler UM, Dunne J and Ulanowicz RE (eds.) Aquatic Food Webs, an Ecosystem Approach, pp. 143--169. Oxford, UK: Oxford University Press. Marshall NB (1971) Explorations in the Life of Fishes. Cambridge, MA: Harvard University Press.
7
Nixon M and Messenger JB (eds.) (1977) Symposia of the Zoological Society of London, No. 38: The Biology of Cephalopods. New York: Academic Press. Robison BH (2003) What drives the diel vertical migrations of Antarctic midwater fishes? Journal of the Marine Biological Association of UK 83: 639--642.
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NEPHELOID LAYERS I. N. McCave, University of Cambridge, Cambridge, UK & 2009 Elsevier Ltd. All rights reserved.
Introduction A remarkable feature of the lower water column in most deep parts of the World Ocean is a large increase in light scattering and attenuation conferred by the presence of increased amounts of particulate material. This part of the water column is termed the bottom nepheloid layer (BNL). Another class of nepheloid layers found especially at continental margins are intermediate nepheloid layers (INLs) (Figures 1 and 2). These occur frequently at high levels off the upper continental slope and at the depth of the shelf edge. From here, they spread out across the continental margin. These INLs are similar to the inversions observed in the BNL on some profiles (Figure 1). (The surface nepheloid layer (SNL), not treated here, is simply the upper ocean layer in which particles are produced by biological activity, and which may have material from river plumes close to shore.) The increase in light scattering is perceived relative to minimum values found at mid-water depths of 2000–4000 m (shallower on continental margins). The increased scattering is due to fine particles. This has been determined by particle-size measurements and filtration of seawater with determination of concentration by weight and volume. Most data on the distribution and character of nepheloid layers have been acquired by optical techniques, principally by the Lamont photographic nephelometer and the SeaTech transmissometer, and more recently the WetLabs transmissometer and light scattering sensor (LSS). The optical work has revealed that the BNL is up to 2000-m thick (can be more in trenches) and generally has a basal uniform region, the bottom mixed nepheloid layer (BMNL), corresponding quite closely to the bottom mixed layer defined by uniform potential temperature (Figure 1). Above the BMNL there is a more or less exponential fall-off in intensity of light scattering up to the clear-water minimum marking the top of the BNL. Both bottom and INLs are principally produced by resuspension of bottom sediments. Their distribution indicates the dispersal of resuspended sediment in the ocean basins and is thus a signature of both material and water transport away from boundaries (some of
8
which may be internal such as ridges and seamounts). Most concentrated nepheloid layers occur on the continental shelf, upper slope, or deep continental margin. They indicate the locus of active resuspension and redeposition by strong bottom currents and internal waves.
Optics of Nephelometers: What They ‘See’ Detection of deep-ocean nepheloid layers has been mainly through measurement of light scattering. The Lamont nephelometer has made the largest number of profiles in all oceans but is no longer in use. It used an incandescent bulb as the source and photographic film as the detector of the light scattered from angles between y ¼ 81 and 241 from the forward axis of the light beam. The film was continuously wound on as the instrument was lowered, resulting in an averaging of the received signal over about 25-m depth. The short-lived Geochemical Ocean Section Study (GEOSECS) nephelometer used a red (l ¼ 633 nm) laser source and a photoelectric cell to detect light scattered from y ¼ 3–151 off the axis of the beam. The SeaTech (now WetLabs, Inc.) LSS measures infrared light (880 nm) backscattered (1801) from particles in the sample volume using a solar-blind silicon detector. Because of multiple scattering, nephelometers do not yield precise optical parameters. Optical transmission with a narrow beam can yield the attenuation coefficient (c). The most commonly used SeaTech and WetLabs transmissometers have a red light source (l ¼ 660 or 670 nm) and usually a 0.20– 0.25-m path length. Most of the contribution to the total scattering b comes from near-forward angles (low values of y). Jerlov (1976) shows that, for surface waters, 47% of b occurs between y ¼ 01 and 31, 79% between 01 and 151, and 90% between 01 and 301. The GEOSECS instrument records about 32% of b and the Lamont nephelometer about 16%. The total scattering is given by Mie theory (assumed spherical particles) as a function of particle size d and relative refractive index (relative RI) n, and wavelength of light l. The relative indices of refraction of suspended material are dominated by components with n ¼ 1.05 and 1.15, values probably characteristic of organic and mineral matter, respectively (e.g., RI of seawater 1.34, quartz 1.55, ratio n ¼ 1.15). Particles from clear ocean waters and weak nepheloid layers (concentration Co40 mg m3) tend
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NEPHELOID LAYERS
9
c 1.0
0.8
0.6
0.4
1.80
1.85
41 24
4700
300
88 16
109
4800
200
179
Pressure (dbar)
215
Meters above bottom
157
234
252
4900
100
227
BMNL 5000 2
4
8 μm
16
32
0
Figure 1 Data from the SeaTech transmissometer in the Atlantic showing the BMNL and a nepheloid layer comprising multiple steps in temperature and turbidity. Turbidity is given as c the attenuation coefficient, and y is potential temperature in 1C. Also shown on the right are particle-size spectra determined by Coulter counter. Reproduced from McCave IN (1983) Particulate size spectra, behavior and origin of nepheloid layers over the Nova Scotian Continental Rise. Journal of Geophysical Research 88: 7647–7666.
to have particle-size distributions by volume which are flat, equivalent to k ¼ 3 in a particle number distribution of Junge type, N ¼ Kd k where N is the cumulative number of particles larger than diameter d and K and k are constants. However, this distribution does not appear to be maintained at sizes finer than about 2 mm where k decreases toward 1.5. In concentrated nepheloid layers, this distribution does not occur at all and a peaked distribution with a peak between 3 and 10 mm is encountered.
Morel has calculated scattering according to Mie theory for suspensions with Junge distributions and several indices of refraction. Recalculation into cumulative curves of percentage scattering in Figure 3 illustrates the fact that most recorded scattering is produced by fine particles. The cases shown are for values of k of 2.1, 3.2, and 4.0 and a two-component (peaked) distribution with k ¼ 2.1 up to a ¼ pdn/ l ¼ 32 and k ¼ 4.0 for larger sizes (a ¼ 32 is equivalent to d ¼ 5.6 mm for l ¼ 633 nm). In each case, three
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10
NEPHELOID LAYERS
(a)
0 Knorr 51 sta. 698 Rockall Trough 54° 23.1′ N, 15° 18.7′ W INL.
Depth (m)
1000
2000 BNL 900m
{
INV 2 INV 1
Nephels, arbitrary scale 3000 (b)
3.0
2.8
2.6
3.2
2200 Nephel inversion 2
Higher-salinity core
2300 NEADW
Depth (m)
2400
2500
Nephel inversion 1 Knorr 51 Station 698 Rockall Trough 54° 28.1′ N 15° 18.7′ W 24 Aug. 1975
2600
AABW Source _ High silica_
2700 40
50
60
70
80
15
90
20
25
30
35
Silica (μg at/l)
Ne
Figure 2 (a) Full-depth profile taken in the Rockall Trough with the GEOSECS nephelometer. An INL and two inversions are apparent. (b) Detail of the lower 500 m of the profile in (a) showing the relationship between the nephel (turbidity) inversions and hydrography. Reproduced from McCave IN (1986) Local and global aspects of the bottom nepheloid layers in the world ocean. Netherlands Journal of Sea Research 20: 167–181.
forward-scattering angles are given. It is clear that scattering close to the beam is more sensitive to large particles than that at 201. (At yo0.51 we are essentially dealing with a transmissometer.) In the case of k ¼ 2.1 only about 22% of the scattering is from
smaller sizes (ao32) at y ¼ 20 1. Thus the curves for y ¼ 10 1 and 20 1 are generally representative, and the distribution for both k ¼ 3.2 and the composite case show that 95% of the scattering is by particles o5 mm for l ¼ 633 nm.
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NEPHELOID LAYERS
16
32
64
128 256
4
8
16
11
32
100 k = 3.2
Cumulative percent scattering
k = 2.1
32
32 32
20°
20° 10°
k = 4.0
2°
20° 10°
2°
50
2° 2°
32
k = 2.1 for < 32
20° 10°
k = 4.0 for < 32
10°
0 1
2
4
8
16
32
64
128 256
0.5
1
2
4
8
1
2
4
8
16
32
= Πdnw / Figure 3 Cumulative percentage of scattering calculated from the data of Morel (1973) by McCave (1986). The right-hand case is for a peaked distribution with the peak at a ¼ 32, equivalent to 5.6 mm for l ¼ 633 nm and n ¼ 1.15. Note that material larger than the peak contributes virtually nothing to the scattering in this case. Reproduced from McCave IN (1986) Local and global aspects of the bottom nepheloid layers in the world ocean. Netherlands Journal of Sea Research 20: 167–181.
So it is dominantly the fine fraction of particles in nepheloid layers that is seen and recorded by nephelometers. These particles have very low settling velocities, less than B5 106 m s1. Although larger particles are present, they are rare and their properties and behavior cannot be invoked to explain features of distribution shown by nephelometers. The SeaTech 0.25 m path length transmissometer has been used for most of the modern work on the structure and behavior of nepheloid layers. The transmission T is related to beam attenuation coefficient c over path length l as T ¼ecl. The major control of c is due to particles. Attenuation is due to absorption a and scattering b, thus c ¼ a þ b. The value of c for pure seawater is about 0.36 m1 for this instrument operating at l ¼ 660 nm, and any excess is due to particulate effects. Because scattering very close to the beam is more sensitive to large particles, the transmissometer is more sensitive to larger particles and the nephelometer to smaller ones (Figure 4).
Nepheloid Layer Features The principal features of nepheloid layers that must be accounted for are the facts that the concentration is generally highest close to the bed, decreasing upward, but also that this is not universally the case, because: inversions, upward increases of concentration, are
also found; steeper gradients in concentration as well as inversions are often found at the boundaries of distinct density (temperature and/or salinity) changes, but their frequency decreases upward. The thickness of the BNL is generally in the region of 500–1500 m, and exceptionally up to 2000 m. This is clearly greater than the thickness of the bottom mixed layer (Figures 2 and 4), a fact which rules out the possibility of simple mixing by boundary turbulence being a sufficient mechanism for BNL generation. In several cases, the nepheloid layer is seen to transcend water masses. That is to say, the nepheloid layer shows a general decline in turbidity upward through interleaved water masses of differing sources and temperature/salinity characteristics though there may be a steeper turbidity gradient at the boundary between water masses. The highest suspended sediment concentrations occur in the BMNL in regions of strong bottom currents where they are typically 100–500 mg m3 (this is the same as mg l1). In general, deep western boundary currents and regions of recirculation carry high particulate loads. However, high turbidity is also found beneath regions of high surface eddy kinetic energy (variance of current speed) when located over strong thermohaline bottom currents. High surface eddy kinetic energy is connected with high bottom eddy kinetic energy; thus, intermittent variability, when added to a strong steady
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12
NEPHELOID LAYERS
mixing (which occurs mainly in the BMNL) up to a kilometer above the bed. It is not possible to mix sediment across sharp density steps without breaking them down. However, these features are explicable if the layers are recently separated from the bottom. Some layers marked by steps in potential temperature contain excess radon-222 (originating from bottom sediments) with a 3.8 day half-life, suggesting detachment of bottom layers within 2–3 weeks before sampling. It is anticipated that with time these layers become thinner by mixing at their boundaries and by lateral spreading to yield, eventually, a uniform stratification. In this, the upper part of the BNL, sheared-out mixed layers that have, on average, come further from the sloping sides of the basins and from regions with less frequent resuspension, have lower concentrations. The basal layers are on average more recently resuspended and also gain material by fallout from above; thus, there is an overall decreasing particulate concentration, and increasing age upward.
−1
PMC (μg l ) 0
0
50
100
150
200
250
Surface nepheloid layer Attenuation excess Coarser/organicdominated particles
200
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600
Shallow intermediate nepheloid layer
Depth (m)
800
1000
t −1
Clear water (~10 μg l ) Good agreement between instruments
1200
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Deep intermediate nepheloid layer 1800
Bottom nepheloid layer Scattering excess Finer/inorganic dominated particles
2000 26.00
26.25
26.50
26.75
27.00
27.25
27.50
27.75
Decay of Concentration: Aging of Particulate Populations 28.00
3
t (kg m )
Figure 4 Profiles of particulate matter concentration calculated from beam attenuation (solid line), and light scattering (dotted line) against depth, together with the density structure (st) of the water column (dashed line). Reproduced from Hall IR, Schmidt S, McCave IN, and Reyss JL (2000) Particulate matter distribution and Th-234/U-238 disequilibrium along the Northern Iberian Margin: Implications for particulate organic carbon export. DeepSea Research I 47: 557–582.
component, may be responsible for very high current speeds which produce intense sediment resuspension.
Separated Mixed-layer Model Nepheloid layer structure is consistent with a quasivertical transport mechanism involving turbulent mixing in bottom layers of B10–50-m thickness (see Turbulence in the Benthic Boundary Layer), followed by their detachment and lateral advection along isopycnal (equal density) surfaces. The detachment occurs in areas of steep topography as well as in areas of lower gradient at benthic fronts where sloping isopycnals intersect the bottom. In many nepheloid layers, there are sharp upward increases in sediment content associated with steps in other properties such as temperature and salinity (Figure 1). Both the step structure and inversions in particulate matter concentration are incompatible with vertical turbulent
The particles composing the nepheloid layer may be modified due to aggregation with similar-sized particles and scavenging by larger rapidly settling ones. Aggregation may also be caused through biological activity, although little is known about such processes at great depths. Particles tend to settle and to be deposited onto the bed, from the bottom mixed layer. The larger particles should be deposited in a few weeks to months, 10–20 mm particles taking 50– 20 days to settle from a 60-m-thick layer. This will not affect the layer perceived by nephelometers so quickly because the timescale of fine particle removal initially involves Brownian aggregation with a ‘halflife’ of several months to years. The direct rate of deposition of very fine particles (0.5–1 mm) from a layer which remained in contact with the bed would be very slow. Concentration would halve in about 8 years. Thus the rate of decrease in concentration of 0.5–1-mm particles is due more to their being moved to another part of the size spectrum by aggregation (and then deposited) than to their being deposited directly. The fine material in dilute nepheloid layers has a mean residence time measured in years, demonstrated by the residence time of particle-reactive short half-life radionuclides such as 210Pb (t1/2 ¼ 22.3 years). In more concentrated nepheloid layers, a large proportion of this material will be removed in under a year, and in the BMNL residence times are tens to a hundred or so days estimated via 234Th (t1/2 ¼ 24.1 days). The
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NEPHELOID LAYERS
dilute nepheloid layers in tranquil parts of the oceans could thus contain material that was resuspended very far away. The contribution of this material to the net sedimentation rate of these tranquil regions may not be negligible. The rate of deposition in the central South Pacific of only 0.5–2 mm ky1 could include up to 1 mm ky1 of fine material from the nepheloid layer. With aging, the individual detached layers comprising the nepheloid layer lose material by aggregation and settling and lose their identity by being thinned through shearing. An originally discontinuous vertical profile of concentration with inversions is converted to one of relatively smooth upward decline
60°
40° E
in concentration. Present understanding of particle aggregation and sinking rates suggests that this takes a few years to achieve.
Chemical Scavenging by Particles in Nepheloid Layers Many chemical species are particle-reactive and rapidly become adsorbed onto surfaces. This is why a number of elements are present in only trace quantities in seawater as outlined by Robert Anderson in 2004. The phenomenon of ‘boundary scavenging’, preferential removal of particle active species at
100°
80°
120° Mid-ocean and aseismic ridges and plateaus
0.2
0°
0.
0.2
E ˚
2
20° N
Excess turbidity [log(Eb /Ec)] values: <0.2
0.2
0.4
0.2_ 0.4 0.4_ 0.6 0.6 _ 1.0
20°
>1.0
0.2
2
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0.
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0.
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0.6 1.0
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1.0
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60° S 0.8
4
0.
0.6
0.4 0.6
0.
6
0.4 0.4 0.8
Figure 5 Distribution of excess turbidity for the Indian Ocean expressed as log(E/Ec) where E is the maximum light scattering near the bed and Ec is the value at the clear water minimum. A value of 1 thus represents a factor of 10 increase from the clear-water value. Reproduced from McCave IN (1986) Local and global aspects of the bottom nepheloid layers in the world ocean. Netherlands Journal of Sea Research 20: 167–181; based on Lamont nephelometer data presented by Kolla V, Sullivan L, Streeter SS, and Langseth MG (1976) Spreading of Antarctic bottom water and its effects on the floor of the Indian Ocean inferred from bottom water potential temperature, turbidity and sea-floor photography. Marine Geology 21: 171–189; and Kolla V, Henderson L, Sullivan L, and Biscaye PE (1978) Recent sedimentation in the southeast Indian Ocean with special reference to the effects of Antarctic bottom Water circulation. Marine Geology 27: 1–17.
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NEPHELOID LAYERS
scavenging in the BNL on the nearby continental margin. The 234Tht (‘t’ means total, i.e., dissolved and particle-associated) profile showed a 234Tht/238U disequilibrium (the difference between 234Tht and 238 U activity) up to at least 500 mab (meters above bottom), and that secular 234Tht/238U equilibrium was reached between 500 and 1000 mab. That is to say radiochemical evidence suggests lateral advection on the 100-day timescale (4–5 234Th half-lives) below 500 mab at a site quite close to the continental margin (about 250 km away). Above that, the timescale is longer.
The Turbidity Minimum The source of nepheloid layers at heights over a few hundred meters above the bed is believed to be the sides of the ocean basins and protrusions from the bottom of the oceans such as seamounts, ridges, and rises. Major source regions are (1) deep zones of stronger currents, which in most areas are at depths greater than 3000 m; and (2) shallow areas where material resuspended at the shelf edge and upper slope by surface and internal wave motions, spreads
15
seaward. Tidal motion on the rough topography of mid-ocean ridges is known to yield enhanced mixing as explained by Kurt Polzin and colleagues in 1997, and is also probably responsible for resuspension. Between these zones are depths of 1000–3000 m where stratification and currents are weak, there is no primary production of organic matter, and the ocean is often undersaturated with respect to calcite, aragonite, and opal. There is thus little particle supply from the side, and a decrease in downward transport due to dissolution and bacterial consumption of CaCO3, SiO2, and organic carbon, leading to a nepheloid minimum where (in the Atlantic) concentrations are 5–20 mg m3. The concentrations given by nephelometers at the clear-water minimum differ by a factor of about 3 between areas under high surface productivity (high) and mid-gyre regions (low). The flux of large biogenic particles from the surface in these areas appears to provide a net supply of smaller particles to mid-depth (rather than scavenging them). This is also seen to be the case in temporal variability at a point, such that higher mid-water turbidity occurs in summer under higher productivity, and is lower in winter.
71° W 22° N
59° W 17° N 2.2°
4
= 2.0°
1.8° 5
6
Caicos outer ridge log E/ED 0.7_ 0.8
7
0.6_ 0.7 0.5_ 0.6
8
Puerto Rico trench axis
Depth (km)
1.6°
1.4° South slope of Puerto Rico trench
0.4
Figure 7 Interflow of suspended sediment generated in the Western Boundary Undercurrent, having detached from the bed on Caicos drift, flowing over the Puerto Rico trench. Reproduced from Tucholke BE and Eittreim SL (1974) The western boundary undercurrent as a turbidity maximum over the Puerto Rico Trench. Journal of Geophysical Research 79: 4115–4118.
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NEPHELOID LAYERS
Concentration and Spreading in the Atlantic and Indian Oceans The broad picture of high-speed inflows on the western sides of ocean basins with a distributed return flow proposed by Henry Stommel has been confirmed by many hydrographic studies and current measurements. These broad patterns of the inflow of cold bottom water are very similar to the excess BNL concentration (excess over the clear-water minimum) for the Indian Ocean, and net BNL particulate load (mass per unit area) for the Atlantic based on Lamont nephelometer data (Figures 5 and 6). It is also apparent that the concentration or load of sediment generally declines going northward along the paths of the major inflows of Antarctic bottom water. However, there are several points of very high concentration or load which are not obviously related to likely increases in mean bottom flow velocity, for example, the high in the center of the Argentine Basin (Figure 6). Richardson and colleagues showed in 1993 that these turbidity highs are sometimes caused by intermittently high velocities under regions where high surface eddy kinetic energy was propagated downward, resulting in high abyssal eddy kinetic energy, but at other times and places there was no correlation with current speeds. Thus some uncertainty remains over the role of high abyssal eddy kinetic energy versus locally accelerated thermohaline flow or deep recirculation loops in producing high concentration. Nevertheless, the major feature of concentrated BNLs is that they do delineate the paths of high-velocity bottom currents extremely well and they decline in concentration away from high-velocity boundary regions and away from the energetic Southern Ocean.
Boundary Mixing, Intermediate Nepheloid Layers, and Inversions Steps and inversions are best shown by modern highsampling-rate sensors (Figures 1, 2, and 4), though striking examples may also be found in older data. INLs are most common off the continental slope and submarine canyons. In both cases, focusing of internal waves (see Internal Waves and Internal Tides) or tides is believed to be responsible, due to the amplification of bottom flow velocities at points where the slope of the seabed S matches the wave characteristic slope C. This causes both resuspension and boundary mixing (see Turbulence in the Benthic Boundary Layer and Estimates of Mixing). The suspensions so generated have both fine and coarse particles at the time of generation but as they spread
seaward they rain out the larger ones over the slope, leaving only finer material with slow sinking speeds. Not all intermediate layers detach on slopes. Perhaps the biggest INL in the ocean occurs where the turbid plume of the North Atlantic Western Boundary Undercurrent leaves the bottom at the end of Caicos Drift off the Bahamas at about 5200 m depth and proceeds at that level over the 8000-m-deep Puerto Rico trench (Figure 7). Pronounced inversions are seen at depths between 1 and 4 km along the continental margin close to the Congo (Zaire) River. A comparable situation is found over Nitinat Fan off the NW USA. This area lies off two submarine canyons which are the most
% Transmission 0
62
64
66
62
64
66
1
2
3 Depth (km)
16
4
5
6
7 9° 17.7′ S, 80° 36.1′ W (Peru trench)
8° 20.4′ S, 81°01.8′ W (Peru trench)
Figure 8 Profiles of the nepheloid layer in the Peru trench made with a 1 m path length transmissometer. Reproduced from Pak H, Menzies D, and Zaneveld JRV (1979) Optical and Hydrographical Observations off the Coast of Peru during May– June 1977, Data Report 77, Ref. 79-14, pp. 1–93. Corvallis, OR: Oregon State University.
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NEPHELOID LAYERS
likely sources of the intermediate layers. Thus, canyons may also act as point sources for supply of turbid layers which mix out into the ocean interior.
Trenches and Channels A final example of the influence of the sides of a basin on its nepheloid layers comes from the extremely thick nepheloid layers found in trenches and some deep-sea passages. The nepheloid layer thickness is that of the trench depth plus the thickness of the nepheloid layer over the adjacent seafloor for the Kuril–Japan trench. The result is a layer 2600-m thick. Nepheloid layers 2700-m thick have been recorded in the Peru trench between 81 and 101 S (Figure 8). Flow through Vema Channel between the Argentine and Brazil basins yielded a well-mixed water column in temperature and turbidity with excess radon-222 up to 400 m above the bed. Models of boundary layer development show that this cannot occur by vertical turbulent diffusion. Turbid layers must have spread quickly from the sides to the center of the channel, giving added support to the detached mixed-layer model. The fact that there is more suspended material at depths greater than about 4000 m partly reflects the fact that waters at those depths are in contact with a much greater area of seabed in proportion to their volume than the shallower parts of the oceans. For 4–5 km depth the value is 0.83 km2 km3, whereas for 2–3 km it is 0.11 km2 km3. One might say that deeper waters feel more bed. The global distribution of nepheloid layers is, for instrumental reasons, the global distribution of fine particles (o2 mm), because that is what the instruments that map them ‘see’, but they also contain larger particles which play an important role in sediment dynamics, especially in the lowest 10 m of the BMNL.
See also Dispersion and Diffusion in the Deep Ocean. Floc Layers. Internal Tidal Mixing. Internal Waves. Marine Snow. Particle Aggregation Dynamics. Transmissometry and Nephelometry. Turbulence in the Benthic Boundary Layer.
Further Reading Amin M and Huthnance JM (1999) The pattern of crossslope depositional fluxes. Deep-Sea Research I 46: 1565--1591. Anderson RF (2003) Chemical tracers of particle transport. In: Elderfield H (ed.) Treatise on Geochemistry, Vol. 6:
17
The Oceans and Marine Geochemistry, pp. 247--273. Oxford, UK: Pergamon. Anderson RF, Bacon MP, and Brewer PG (1983) Removal of 230Th and 231Pa at ocean margins. Earth and Planetary Science Letters 66: 73--90. Armi L and D’Asaro E (1980) Flow structures of the benthic ocean. Journal of Geophysical Research 85: 469--484. Bacon MP and Rutgers van der Loeff MM (1989) Removal of thorium-234 by scavenging in the bottom nepheloid layer of the ocean. Earth and Planetary Science Letters 92: 157--164. Baker ET and Lavelle J W (1984) The effect of particle size on the light attenuation coefficient of natural suspensions. Journal of Geophysical Research 89: 8197--8203. Biscaye PE and Eittreim SL (1977) Suspended particulate loads and transports in the nepheloid layer of the abyssal Atlantic Ocean. Marine Geology 23: 155--172. Dickson RR and McCave IN (1986) Nepheloid layers on the continental slope west of Porcupine Bank. Deep-Sea Research 33: 791--818. Gardner WD (1989) Periodic resuspension in Baltimore Canyon by focusing of internal waves. Journal of Geophysical Research 94: 185--194. Hall IR, Schmidt S, McCave IN, and Reyss JL (2000) Particulate matter distribution and Th-234/U-238 disequilibrium along the Northern Iberian Margin: Implications for particulate organic carbon export. Deep-Sea Research I 47: 557--582. Jerlov NG (1976) Marine Optics. Amsterdam: Elsevier. Johnson DA, McDowell SE, Sullivan LG, and Biscaye PE (1976) Abyssal hydrography, nephelometry, currents and benthic boundary layer structure in Vema Channel. Journal of Geophysical Research 81: 5771--5786. Kolla V, Henderson L, Sullivan L, and Biscaye PE (1978) Recent sedimentation in the southeast Indian Ocean with special reference to the effects of Antarctic bottom water circulation. Marine Geology 27: 1--17. Kolla V, Sullivan L, Streeter SS, and Langseth MG (1976) Spreading of Antarctic bottom water and its effects on the floor of the Indian Ocean inferred from bottom water potential temperature, turbidity and sea-floor photography. Marine Geology 21: 171--189. McCave IN (1983) Particulate size spectra, behavior and origin of nepheloid layers over the Nova Scotian continental rise. Journal of Geophysical Research 88: 7647--7666. McCave IN (1986) Local and global aspects of the bottom nepheloid layers in the world ocean. Netherlands Journal of Sea Research 20: 167--181. McPhee-Shaw E (2006) Boundary–interior exchange: Reviewing the idea that internal-wave mixing enhances lateral dispersal near continental margins. Deep-Sea Research II 53: 42--59. Menard HW and Smith SM (1966) Hypsometry of ocean basin provinces. Journal of Geophysical Research 71: 4305--4325. Morel A (1973) Indicatrices de Diffusion Calcule´es par la The´orie de Mie pour les Syste`mes Polydisperses, en Vue de l’Application aux Particules Marines, Report 10, pp.
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18
NEPHELOID LAYERS
1–75. Paris: Laboratoire d’Oce´anographie Physique, University of Paris VI. Pak H, Menzies D, and Zaneveld JRV (1979) Optical and Hydrographical Observations off the Coast of Peru during May–June 1977, Data Report 77, Ref. 79-14, pp. 1–93. Corvallis, OR: Oregon State University. Polzin KL, Toole JM, Ledwell JR, and Schmitt RW (1997) Spatial variability of turbulent mixing in the abyssal ocean. Science 276: 93--96. Richardson MJ, Weatherly GL, and Gardner WD (1993) Benthic storms in the Argentine Basin. Deep-Sea Research 40: 957--987. Thorndike EM (1975) A deep sea, photographic nephelometer. Ocean Engineering 3: 1--15.
Tucholke BE and Eittreim SL (1974) The western boundary undercurrent as a turbidity maximum over the Puerto Rico Trench. Journal of Geophysical Research 79: 4115--4118. Turnewitsch R and Springer BM (2001) Do bottom mixed layers influence 234Th dynamics in the abyssal nearbottom water column? Deep-Sea Research I 48: 1279--1307. Zaneveld JRV, Roach DM, and Pak H (1974) The determination of the index of refraction of oceanic particulates. Journal of Geophysical Research 79: 4091--4095.
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NETWORK ANALYSIS OF FOOD WEBS J. H. Steele, Woods Hole Oceanographic Institution, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1870–1875, & 2001, Elsevier Ltd.
Introduction Photosynthesis transforms energy from sunlight into calories within marine plants, predominantly phytoplankton and seaweeds. The plants use this energy to take up carbon and essential nutrients, such as nitrogen and phosphorus, from sea water to produce organic materials.This organic matter forms the base for the food web composed of herbivores thateat those plants and the carnivores that prey on the herbivores. There can beseveral trophic levels of carnivores, including all the fish species that weharvest. In the open sea, away from the coast and theseabed, microscopic single-celled phytoplankton dominate the plant life, so the organisms tend to get bigger as each trophic level feeds on the onebelow: from small herbivorous crustaceans to larger invertebrates, to small andlarge fish, and finally to human beings and marine mammals. When any animal consumes food, most of the energy in that food is used for metabolism; some of the remainder is excretedas waste products and only a small fraction goes to growth. In young, coldblooded animals in the sea, growth can be relatively efficient: 20–30% of energy intake. In older animals growth is replaced by reproduction, which, after all, is the whole point of the life cycle. As an approximate overall figure we usually assume that the energy converted into growth and reproduction is about 10% of the total energy intake. Thus, in asimple trophic pyramid, the energy in successive trophic levels would be 100:10:1:0.1. From this one can see why we are encouraged to eat plants on land, and why fish from the sea are energetically expensive in terms of plant calories. In practice it is very difficult to measure directly the energy content of marine organisms and, especially the energy transfers between trophic levels. However, carbon is the essential building block for organic matter, being taken up from inorganic form at photosynthesis and returned to sea water during respiration. The carbon content of organisms and the rates of uptake of inorganic carbon can be measured using radioactive carbon, carbon-14, as a tracer;
transfers through the food can then bemeasured. Carbon is therefore frequently used as a proxy for the more elusive concept of energy flow. All organisms also require many essential elements, and many of these are in short supply in sea water. In particular, inorganic nitrogen and phosphorus, as nitrate and phosphate, are regarded as limiting factors in photosynthesis and thus in the rate at which energy and carbon are supplied to the food web. Since organic carbon, nitrogen and phosphorus have a roughly constant ratio in marine organisms (the Redfield ratio), nitrogen can also be used as a proxy for energy flow; it will not, however be considered here. Carbon and nitrogen fluxes are also important in relation to other issues concerning marine food webs. The biologically mediated flux of carbon to deeper water is important for the calculation of global carbon budgets and climate change. Eutrophication in coastal waters produces imbalance in the food webs of these regions.
Units Ideally, all presentations would be in units of energy but, as discussed, the actual measurements are often in carbon or, for fisheries, in biomass (wet weight). Conversion from one unit to another will vary with the organism but, as an approximation: 10 kcal ¼ 1g carbon ¼ 10 g biomass:
½1
History The first attempt to produce an energy budget for an aquatic food web was made by Lindeman in 1942, for a freshwater lake. In 1965 a budget for the North Sea (Figure 1), showed possible pathways from primary production to commercial fisheries. The main conclusion was that the ecological efficiencies needed to be quite high for the budget to balance. This conclusion depends on the number of pathways introduced; for example, the category ‘other carnivores’ is put in to account for the presence of gelatinous invertebrate predators, such as ctenophores, that are ubiquitous but the biomass of which is difficult to estimate. Without this box, the herbivore efficiency could be less than 20%. In 1969 Ryther wrote a seminal paper that outlined the pathways at the global scale, by dividing the marine ecosystems into three types, upwelling,
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19
20
NETWORK ANALYSIS OF FOOD WEBS
Net primary production
Table 1 Primary production, number of trophic levels to commercial fish species, efficiency of transfer of energy (or carbon) between trophic levels and resultant ratios of fish to primary production for three major marine categories
(Detritus) Meiobenthos
Macrobenthos
Other carnivores
Demersal fish
Pelagic herbivores Other carnivores
Pelagic fish
Primary production (gC m2 year1) Trophic levels Ecological efficiency (%) Fish production: primary production (%)
Large fish Yield to humans
(A)
100 20
Coastal
Upwelling
50
100
300
5 10 0.01
3 15 0.3
1.5 20 12
80
8 2
Ocean
12
1 2
16
3
4
0.4
0.17
0.3
12
0.6
1.8
0.04
gelatinous predators, there is still relatively little quantitative information on biomass and production; this has led to an extensive development of mathematical treatments to infer energy or carbon flows.
0.04 (B)
0.13
0.3
25%
23%
20%
13%
10%
10%
Quantitative Methods
15%
(C) Figure 1 Energy or carbon fluxes in the food web for the North Sea.(A) A simplified food web. (B) Transfers of particulate organic carbon assuming a primary production of 100 gC m2 peryear. Each box indicates the production by that component available to its predators, including humans. (C) The transfer or ecological efficiencies calculated from the output as percentage of input.
continental shelf, and open ocean(Table 1). The major implication of Ryther’s calculations was that there were no great untapped fish resources in the open ocean. He estimated that the potential sustained yield offish to humans was unlikely to be greater than 100 million tons. In 1998, the yield was about 90 million tons. The reason for the low open-oceanyield of fish – and the controversial part of Ryther’s calculation– is his estimate of five trophic levels from primary production to fish. This choice was based on the very small size of open-ocean phytoplankton. In the intervening years, research on the base of open-ocean food webs has shown that much of the primary production is recycled through the smallest sized categories in the food web – the microbial loop. For the intermediate trophic levels, such as the
Early calculations, such as those illustrated in Figure 1, were put together by a very informal series of iterations until all the flows balanced and the ecological efficiencies were not unreasonable. The major advance has been the development of numerical methods to make more objective estimates of best fit. Essentially, the ecosystem is considered to be at steady state, so that, for any box, such as those in Figure 1, there has to be a balance between rates of energy flow or carbon flux entering and leaving: consumption ¼ predation þ metabolism þ exportðinputÞ
½2
The information for each of these terms for each box canbe qualitatively different. In particular, for animals: consumption ¼ growth þ reproduction þ metabolism ½3 These complexities are eliminated if it is assumed that, for each box, there is a constant ecological efficiency, ei : ei ¼
ðgrowth þ reproductionÞ consumption
½4
Generally, for the system to be soluble, the terms in eqn [2] for all boxes must be expressed as a set of linear equations. For this purpose the variables are usually taken to be the energy, carbon or biomass flux through each box, Ti , that is available to higher
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NETWORK ANALYSIS OF FOOD WEBS
trophic levels. Linearity requiresthat the rate processes in eqn [2] are constants independent of Ti . Ti X ¼ bij Tj þ cj i ¼ 1; n; ½5 ei where bij is the fraction of Tj consumed by Ti and ci is the constant rate of external input to Ti ; ci would be, for example, the rate of primary production. Note that, as in Figure 1, all the variables are rates of throughput in the food web and the parameters are non-dimensional.These matrix inversion techniques can also be used with nutrient flows that involve recycling. An alternative top-down approach is often used for flow calculations where emphasis is on the higher trophic levels, and fish yields are the defining input. For these situations,consumption, Ci , is expressed as: ðconsumptionÞ biomass biomass Ci Bi ¼ Bi
consumption ¼
½6
Then the biomass in each box, Bi , becomes the state variable. The rate process Ci =Bi is assumed constant, for each box. Then: ai Bi ¼
X
pij Bj þ di ;
½7
where ai represents production per unit biomass, the ‘P/B’ ratio, assumed constant for any box; pij is the unit consumption rate of Bj on Bi , assumed to be independent of the magnitude of Bj or Bi ; di is the export, assumed constant for each box. If all the parameters are known for either eqns [2] and [7] then the set of equations canbe solved for Ti or Bi. In practice it is never as simple as that. Usually a number of parameters are unknown or ill defined; for example, as upper or lower bounds.Then iterative procedures can be used to obtain best fits. The selection of the number of boxes and the content of each box is the critical process. The assumption of a linear, steady-state ecosystem is the critical constraint for this type of analysis involving a large number of boxes. This approach is complemented by the highly nonlinear analysis used for modelling offisheries and plankton dynamics.
Examples of Carbon and Biomass Networks The emergence of the microbial loop as asignificant feature of pelagic food webs, together with the availability of computer based inverse methods, resulted in a focus on flow analysis of the lower levels
21
of the pelagic ecosystem; levels that were represented simply as phytoplankton-to-zooplankton in Figure 1. The results of analyses (Figure 2) illustrate how this part of the system isexpanded into seven boxes with 19 links. Two examples in Figure 2, from the continental shelf around Britain, show the kinds of patterns that arise from these calculations. The authors point out differences between the English Channel and the Celtic Sea. The former puts 85% of primary production through the microbial loop, whereas the latter has 40% going directly to the mesoplankton (predominantly copepods). However,for both examples, the exports from this part of the food web are very similar: 3% to higher trophic levels via the mesoplankton, and 30% to the benthos as detritus. The major change from the earlier calculations is the dominant role of the microbial loop, involving recycling of most of the primary production.Only the ‘new’ production from nitrate (NO3) in Figure 2A fuels the export of detritusand mesoplankton to higher trophic levels. The alternative approach (eqn [7]) has been used for a wide range of aquatic ecosystems, including lakes and coral reefs as well as ocean systems. One example (Figure 3) shows calculations for the continental shelf around the Gulf of Mexico. There is no microzooplankton box and the general flow patterns are not dissimilar to those of Figure 1; 30% of primary production goesdirectly to detritus, as does 40% of zooplankton food, presumably as fecal material. Detritus then feeds the demersal fish through the benthos. Note, however, that once again detritus is the dominant biomass (85% of the total). It would appear that all the solutions of these linear systems – whether as energy or biomass, for lower or higher trophic levels – require a major role for detritus. Detritus is a difficult variable to define and measure. This box can contain phytoplankton cells, zooplankton feces, marine snow and other residues. It is assumed that detritus is broken down by bacteria but the ai or ei values are not well known. It is rarely sampled directly, eitheras biomass or for rate processes, and so is an empty box that can be used to balance flows.
Discussion In terrestrial ecosystems most of our food comes directly from plants, with the remainder from herbivores – animalsthat eat plants. In the sea, the fish we eat are nearly all carnivores. Many feed on the small herbivorous crustaceans, such as krill, but some – the most highly prized, like tuna – themselves feed on carnivores, such as smaller fish. So what we hope to take from the sea, the potential fishery yield,
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22
NETWORK ANALYSIS OF FOOD WEBS
Sea surface A
NH4
Mi
NO3
B
Me
DOM
P Export
D Euphotic zone Aphotic zone
(A)
Sea surface GPP
A 169
720
71
257 91
Me 95
28 104
Mi 126
160 31.5 41.5
41.5 107 9.6
25.4 141
36 11.4
P 46
B 35.8 DOM 41.9 194 15.1
Export D
Euphotic zone Aphotic zone
216
(B)
Sea surface GPP
A 71.9
729
72
163 285 39 99
Me 227 23 136
Mi 67.9
22.4
103
62
75 73
60
P 31.4 51.8 15.8 4.1
B 23.4 DOM 60.5 210 17.1
X dBi ¼ ai Bi pij Bj di dt
Export D 227
(C)
depends very much on the patterns and magnitudes of the flow of energyor carbon from primary production. Our understanding of these patterns cantherefore provide valuable estimates on the limits of what we can take from thesea, as well as increasing our insight into the ecological processes. The selection of boxes and the arrows between them depends on our knowledge of which prey– predator interactions are quantitatively significant. It is obvious from the examples here that thereis a considerable compaction into large, often heterogeneous groups, such as mesoplankton or pelagic fish. Thus the level of organization is very differentfrom that for biodiversity studies, or even for descriptions of the full complexity of the food web. However, in producing a manageable number of boxes, it is important not to fold significant prey– predator interactions into the same box. Thus, in Figure 1, prey–predator interactions at the microbial scale were ignored. On the other hand, it has been pointed out that, for the North Sea, the inclusion of both the detrital box and the invertebrate carnivores does not provide enough energy flow for the fishery yield. The definition of boxes will remain acentral problem, but this is also the great strength of this approach: it requires attention to all aspects of the food web. The major limitation is that the method assumes that the system is in steady state. This is usually achieved by taking yearly averages and ignoring shorter seasonal variability and longer decadaltrends. For the former, several researchers have constructed dynamic planktonmodels of the nonlinear interactions between nutrients and detritus based onthe seven boxes or variables in Figure 2. The longer term changes are especially important for fisheries. It is possible to transform the linearized eqn [7] for equilibrium states into a time varying system by writing:
Euphotic zone Aphotic zone
Figure 2 (A) Generic model of a plankton food web in the upper layer of a stratified water column. (B) The inverse solution for carbon fluxes in summer at a station in the English Channel. Values inside the boxes are respiration flows (mgC m2 per day).(C) Flows at a station in the Celtic Sea. A, autotrophs; B, bacteria; D, detritus; DOM, dissolved organic matter; solid arrows denote intercompartmental transfer of carbon; dashed arrows indicate flows of inorganic nitrogen into the system; Me, mesozooplankton; Mi, microzooplankton; P, Protozoans. (Vezina and Platt, 1988.)
½8
This linearized approximation to an essentially nonlinear system can be used to indicate the directions that changes may take when the system departs from a previous steady state, but is unlikely to be adequate for the very large switches, or regime shifts, that occur in the relative abundance of different fish stocks. An alternative is to assign very different values to the boxes for the fish stock biomass and then recalculate the network. This technique can be used to estimate the status of ecosystems before the impact of human predation on fish or marine mammals. As an example,the Gulf of Mexico flows were
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(c) 2011 Elsevier Inc. All Rights Reserved.
1
2
3
444.8
17.3
0.1
311.7
1776.51
Detritus B = 236.0
200
140.01
0.2
B = 0.13
54.1
0.7 1.3 209.7
0.03 0
10031
0.01
11.51
Pelagic predators B = 0.05
0.0 0.02 0.3
Phytoplankton B = 4.7
5420
1970.61
B = 3.6
391.4
06 0.01 13.5 0.4 213.7 313.6 Zooplankton 28.5
256.8
0.01 0
0
Billfish B = 0.01
0.7
10.81
Dolphin B = 0.02
0.1
1.5
0.03
1270.31
Pelagic fish B = 12.5
0.2 2.7 0.02
10.71
0.1
2.7 Mackerels 0
0.7
0.02 13.31 0
B = 0.04
0 0.5 Tunas
4.4
00.4
0
Macro- 9.8 crustaceans B = 1.0
0.4
8.0
0.03
Sharks B = 0.08
10.71
0.1
0.2
18.71
B = 0.7
2.7
0.8 0.2
40.4
2.0 20
Benthos plants B = 3.7
18.3
133.01
6.5
Benthos B = 5.0
88.7
11
36.7
18.33
36.2
10.9
153.51
Demersal fish B = 3.5
17.6
0.6 0.4 4.4
0.2 0.04 1.6 0.04 Demersal 6.4 0.3 predators
Other exports
Respiration
Harvest
Flow to detritus
Figure 3 Gulf of Mexico ecosystem model: trophic compartments with values for biomass (g m2) and connecting flows (gm2 per year1). (Christensen and Pauly, 1993.)
Trophic level
4
NETWORK ANALYSIS OF FOOD WEBS 23
24
NETWORK ANALYSIS OF FOOD WEBS
recalculated, with the biomass of the top predators increased by a factor of 10. The major change was that the utilization of detritus within the ecosystem increased from 11% to 70%. Thus the ‘detritus’ box appears to act as a reservoir for over-exploited systems.
Conclusions Calculations of the overall energy, orcarbon, fluxes through a marine ecosystem provide a valuable check on estimates for the potential productivity associated with each component of the food web. These calculations can set limits on expected yields to humans; they can act as links between detailed, but necessarily static, descriptions of biodiversity, and models of the dynamics of individual populations. It is necessary to bear in mind that our knowledge of the food webs used in these calculations are provisional and the outcome of the calculations is dependent on the specification of this structure; thus the full benefits from this approach depend on further studies of each ecosystem.
Further Reading Christensen V and Pauly D (eds) (1993) Trophic models of aquatic ecosystems. ICLARM Conf. Proc. 26: 390pp. Christensen V and Pauly D (1998) Changes in models of aquatic ecosystems approaching carrying capacity. In: Ecosystems management for sustainable marine fisheries. Ecol. Appl 8 (suppl). Fasham MJR (1985) Flow analysis of materials in the marine euphotic zone. In: Ulanowicz RE and Platt T (eds) Ecosystem theory for biological oceanography. Can. Bull. Fish Aquat. Sci. 213: 260pp. Lindeman RL (1942) The trophic-dynamic aspect of ecology. Ecology 23: 399--418. Ryther JH (1969) Photosynthesis and fish production in the sea. Science 166: 72--76. Steele JH (1974) The Structure of Marine Ecosystems. Boston: Harvard University Press. Vezina AF and Platt T (1988) Food web dynamics in the ocean. I. Best-estimates of flow networks using inverse methods. Mar. Ecol. Prog. Ser 42: 269--287. Wulff F, Field JG, and Mann KH (eds.) (1989) Network Analysis in Marine Ecology. Berlin: Springer-Verlag.
See also Carbon Cycle. Copepods. Diversity of Marine Species. Eutrophication. Krill. Large Marine Ecosystems. Marine Snow. Microbial Loops. Photochemical Processes. Upwelling Ecosystems.
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NEUTRAL SURFACES AND THE EQUATION OF STATE T. J. McDougall and D. R. Jackett, CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia & 2009 Elsevier Ltd. All rights reserved.
Introduction The equation of state of seawater relates the in situ density r of a parcel of seawater to its salinity S, temperature t, and pressure p. Because in situ temperature t is not a conservative variable, it is often more convenient to express the equation of state in terms of the potential temperature y: potential temperature is the temperature that a fluid parcel would have if its pressure were changed adiabatically and without exchange of matter, from the in situ pressure p to a fixed reference pressure pr . The equation of state then takes the functional form r ¼ r(S, y, p). Even without the exchange of heat and salt so that a fluid parcel retains its salinity and potential temperature, the fact that seawater is compressible means that the parcel’s in situ density r will change as it moves around the ocean experiencing different pressures. In order to counteract the effect of compressibility, Wu¨st in 1933 introduced ‘potential density’ which is the density that a fluid parcel would have if its pressure were changed adiabatically and without exchange of salt to an arbitrarily chosen, but fixed, reference pressure. Oceanographers often use sea surface pressure as the reference pressure, pr ¼ 0 and potential density is often given the symbol ry ¼ r(S, y, 0). However, because the compressibility of seawater is not constant, when a fluid parcel is moved a small distance along a potential density surface, a vertical force must be supplied to keep the parcel on that surface; in other words, a potential density surface is not a surface of neutral or zero buoyancy. Because of the theoretical difficulties in forming a neutral surface throughout extended regions of the ocean, we initially restrict attention to the requirement for a tangent plane to possess the property of neutral buoyancy. A neutral tangent plane in threedimensional space is defined so that an in situ parcel of seawater can be moved small distances along the neutral tangent plane without experiencing a vertical buoyant restoring force. The partial derivatives of
density are used to define the three coefficients 1 @r 1 @r ; a¼ ; b¼ r @Sy;p r @yS;p
and
1 @r k¼ ½1 r @pS;y
where b is the saline contraction coefficient, a is the thermal expansion coefficient, and k is the adiabatic (and isentropic) compressibility (rk ¼ c2 where c is the speed of sound in seawater). When a fluid parcel is moved a small distance adiabatically and without change in salinity in the neutral tangent plane, neither the potential temperature nor the salinity of the parcel changes during this movement, so the parcel’s in situ density changes only in response to its change in pressure, multiplied by the compressibility k. The definition of the neutral tangent plane in terms of the lack of buoyant restoring forces implies that the parcel’s in situ density must be the same as that of the seawater that surrounds the parcel. In this way, the two-dimensional gradients of in situ density and pressure in the neutral tangent plane obey rn ln r ¼ krnp, and expanding the gradient of in situ density in terms of the gradients of salinity, potential temperature, and pressure using eqn [1], we find that the spatial property gradients in the neutral tangent plane obey both rn ln r ¼ krn p
and
brn S ¼ arn y
½2
Here rn is the two-dimensional spatial gradient operator for properties measured in the sloping neutral tangent plane although the horizontal elements of distance are measured in the geopotential surface. This definition of the neutral tangent plane in terms of the lack of buoyant restoring forces implies that the normal n to the neutral tangent plane is in the direction ary brS ¼ krp rln r
½3
which is written in terms of three-dimensional spatial property gradients. Oceanographers care about neutral tangent planes because the turbulent diffusivity that represents vertical (or diapycnal) diffusion is approximately 8 orders of magnitude smaller than the diffusivity that represents the turbulent mixing of properties along neutral tangent planes. In order to separately understand the very different lateral and vertical mixing processes it is necessary to be able to
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25
26
NEUTRAL SURFACES AND THE EQUATION OF STATE
determine the slope of the relevant mixing directions (i.e., the slope of the neutral tangent plane) to an accuracy of about 10 4. As explained by McDougall and Jackett, it is the smallness of the rate of dissipation of mechanical energy in the ocean that confirms the common assumption that the strong lateral diffusion achieved by mesoscale eddies occurs along neutral tangent planes in the ocean interior.
Requirements for a Neutral Surface to Exist Intuitively one might expect that since neutral tangent planes exist everywhere in the ocean, these planes would join to form a series of ‘neutral surfaces’. However, it is well known that surfaces normal to the vector n are only well defined if n r n is everywhere zero. Since the normal to the neural tangent plane is in the direction ary brS the requirement that a neutral surface exists in some region of physical space is that neutral helicity defined by H ðary brSÞ r ðary brSÞ
½4
is zero everywhere in that region. McDougall and Jackett have shown that neutral helicity can be written as H ¼ bTb rp rS ry
½5
where the thermobaric parameter Tb which represents the key nonlinearity of density, is Tb ¼ ap ða=bÞbp ¼ bða=bÞp
½6
From figure 9(b) of McDougall’s 1987 article (featured in the Journal of Geophysical Research) we see
that a value of TbE2.7 108 K1 db1 ¼ 2.7 1012 K 1 Pa 1 is typical of the bulk of the ocean’s volume. Equation [5] says that if neutral helicity is zero, the line of intersection of the S and y planes, rS ry, must lie in the isobaric surface; approximately, this means that rS ry must be a horizontal vector. A sketch of the S, y, r, and p planes and the lines rS ry and rp rr is shown in Figure 1. The neutral tangent plane is always well defined and it includes both the lines ry rS and rp rr, but a neutral surface is well defined only if these two lines coincide. On the face of it, this requirement seems very restrictive; one would think that the distributions of salinity and temperature would demand more freedom than to be bound by this mutual constraint. Nevertheless, McDougall and Jackett showed that the ocean’s hydrography is close to obeying this seemingly unreasonable constraint, implying that all six colored planes of Figure 1 intersect along the single line ry rS//rp rr. Because neutral helicity is found to be quite small in the ocean one is able to construct properly defined surfaces whose tangent planes give very good approximations to the neutral tangent planes. One way of forming such approximately neutral surfaces is described in the section titled ‘Neutral density surfaces compared with potential density surfaces’.
The Helical Nature of Neutral Trajectories When one forms an approximately neutral surface (such as can be found by using the Neutral Density software of Jackett and McDougall), one indeed
Neutral tangent plane
∇
×∇ S
S
∇p ×
∇
Figure 1 Sketch of the planes of constant salinity S, potential temperature y, potential density ry, all intersecting along the line ry rS and the planes of constant pressure p, and in situ density r, that intersect along the line rp rr The neutral tangent plane includes both the lines ry rS and rp rr. A neutral surface exists if and only if these two lines coincide so that the line ry rS lies in the isobaric surface. From McDougall TJ (1988) Neutral-surface potential vorticity Progress in Oceanography 20: 185–221.
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NEUTRAL SURFACES AND THE EQUATION OF STATE
27
urface
Neutral helix
Sea s
b Approximately neutral surface
tant p cons
a e θ,S
co
, S
co
ns
ta
nt
ns
tan
t
c d
nt
p consta
z y Bottom
x Figure 2 The neutral helix depicts circulation such as would occur in diapycnal mixing. The helical nature means that mean vertical motion surfaces.
a trajectory of the mean the absence of interior of this neutral trajectory occurs through density
finds a mathematically well-defined surface but such a surface will not exactly obey the neutral property that fluid parcels on the surface can be moved small distances in the surface without feeling buoyant restoring forces. This is illustrated in Figure 2, which shows schematically a spiraling motion along a sequence of neutral directions following the mean circulation. Even in the absence of small-scale diapycnal mixing processes, flow along this neutral helix will achieve mean dia-surface transport though any real density surface that one cares to construct. This spiraling motion amounts to a new diapycnal advection mechanism that operates independently of the diapycnal motion caused by breaking internal waves. A present challenge is to properly quantify this upwelling mechanism. Further insight into the cause of the helical nature of neutral trajectories can be gleaned from Figure 3. The first leg of the neutral trajectory (from a to b) occurs at constant pressure with the salinity and temperature decreasing in a neutral manner according to eqn [2]. The next leg has the salinity and temperature constant as the pressure increases; the third leg has increasing temperature and salinity (in accordance with eqn [2]) at constant pressure. The fourth leg (from d to e) has constant temperature and salinity while pressure decreases. The reason why the final point along the neutral trajectory, e, does not return to point a can be understood from the Sy diagram of Figure 3. The changes in salinity and potential temperature that occur along legs ab and cd occur at different pressures, so these legs appear in the Sy diagram of Figure 3 as potential density
a Slope = (pa)
d,e
Slope = z /S z
b,c
Slope = (pc)
S Figure 3 An idealized example of how mean vertical advection arises from the ill-defined nature of neutral surfaces. From McDougall TJ and Jackett DR (1988) On the helical nature of neutral trajectories in the ocean. Progress in Oceanography, 20, 153–183.
isolines, but they are potential densities referenced to different pressures. These lines have different slopes because the ratio of the expansion coefficients, b/a, is a function of pressure.
Neutral Density Surfaces Compared with Potential Density Surfaces This article has so far concentrated on the theoretical reasons why neutral surfaces do not exist in the ocean. The field of oceanography has used several different types of surfaces that to varying degrees approximately possess the neutral property. The first and most obvious surface to consider is a surface of constant in situ density. It is easy to demonstrate that seawater is sufficiently compressible in comparison
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NEUTRAL SURFACES AND THE EQUATION OF STATE
with the vertical stratification of the ocean so that an in situ density surface is often quite close to being isobaric. In fact, sound fixing and ranging (SOFAR) floats are much less compressible than seawater and thus these floats have constant mass and almost constant volume so that their density is almost constant. Hence these floats move around the ocean close to in situ density surfaces, and these floats are often described as ‘isobaric floats’. Potential density surfaces have been the most commonly used approximation to neutral mixing directions for much of last century. Potential density is defined to be the density that a parcel of seawater would have if its pressure were changed in an adiabatic and isohaline fashion from its in situ pressure to a fixed reference pressure. Typically reference pressures of 0, 2000, and 4000 dbar are chosen. In the last half of the nineteenth century, Reid and Lynn improved on the straightforward use of potential density by linking successive potential density surfaces that were referenced to different reference pressures. The choices in their method were made based on latitude, and importantly, the particular potential density surface that spanned the equator was common to the definitions that applied to the Northern and Southern Hemispheres. They showed that at the sea surface outcropping of their patched potential density surfaces, the in situ density was different by up to 0.14 kg m3 between the Northern and Southern Hemispheres. This offset of both in situ density and potential density between the hemispheres must be respected by any variable or surface that approximates the neutral property. By so doing, such a ‘neutral density’ variable becomes nonthermodynamic, since it depends on not only salinity, temperature, and pressure, but also location. Jackett and McDougall published a computer algorithm in 1997 for determining a new density variable called neutral density gn for ocean data. Underlying this algorithm is an accurately labeled world ocean data set that has been prelabeled with the neutral density variable. A gn label for an arbitrary {S, T, p, lat, long} ocean observation is determined by first finding the depths on the surrounding four casts in the reference data set where the hydrographic properties of the {S, T, p} observation and the four points in the reference data set are on the same neutral tangent plane. The gn label of the {S, T, p, lat, long} observation is then taken as a weighted average of the interpolated gn values in the prelabeled data set. For the global neutral density variable gn defined by Jackett and McDougall, the global reference data set was based on the climatological atlas of Levitus, with slight modifications having been made to account for extremities in
density and an adequate resolution of the Antarctic Circumpolar Current. Since the normal n to the neutral tangent plane is in the direction of brS ary, the variable gn in the reference data set is taken to approximately satisfy rgn ¼ brðbrS aryÞ
½7
where b is an unknown scalar function of space. This equation is a system of first-order hyperbolic partial differential equations, generally solved using the method of characteristics. The characteristic directions of eqn [7] turn out to be the neutral tangent planes which satisfy brnS ¼ arny and along which gn is approximately constant, so the method of characteristics for this problem reduces to fitting approximately neutral surfaces to the three-dimensional reference data set, with the values of gn on these surfaces being determined by a knowledge of gn down just one cast. The noncharacteristic cast we choose in our global data set is located at 161 S, 1881 E, down which we defined gnref ðzÞ ¼ ry ð0Þ 1000 kg m3
Z0
rg1 N 2 dz
½8
z
being a simple integration of eqn [7] in the vertical with a b profile of unity. The equation of state used for the calculations of r, a, b, and N2 in eqns [7] and [8], is the UNESCO equation (see Appendix 5. Properties of Seawater). The ambiguity in defining neutral density manifests itself in that gn is not exactly constant on local neutral tangent planes. When one integrates along a neutral trajectory (which everywhere is parallel to neutral tangent planes) around loops of global scale one finds that the vertical excursion of such a helix is of the order of tens of meters, consistent with the small size of neutral helicity eqn [5]. The final important stage in generating the gn field in the global reference data set consists of averaging this error over the entire globe. This was achieved by iteratively averaging to a steady state the values of the initial gn field on local tangent planes calculated in the northward, southward, eastward, and westward directions at each location: in effect, a cellular automata of the world ocean. Notice that the initial value of b ¼ 1 down the Pacific Ocean reference cast in eqn [8] was slightly altered during this averaging process. Figure 4 shows various density surfaces along the 3281 E longitude meridian in the Atlantic Ocean taken from the Koltermann et al. hydrographic atlas of the world ocean. The red line shows a neutral trajectory from south to north calculated by joining accurately defined neutral tangent planes in this
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NEUTRAL SURFACES AND THE EQUATION OF STATE
29
0
Pressure (dbar)
1000
2000
3000
4000
5000
6000 −80
−60
−40
−20
0 Latitude
20
40
60
Figure 4 Surfaces of constant gn (blue solid line), sy and s2 (dashed lines with s2 being the lower dashed line), and a neutral trajectory (red line) along 3281 E in the Atlantic Ocean. All the surfaces coincide at 250 dbar and 591 S.
north–south direction, while the blue line shows the gn surface obtained by interpolating the gn field after labeling with neutral density. The two surfaces coincide at 591 S at 250 dbar pressure. The two dashed lines are potential density surfaces referred to pressures of 0 dbar (sy) and 2000 dbar (s2), again coinciding with the neutral trajectory at 591 S. The figure demonstrates the dramatic improvement in the accuracy of neutral density gn in approximating the neutral trajectory over the more traditional potential density surfaces. The maximum pressure difference between the gn surface and the neutral trajectory is 50 dbar while the maximum pressure difference for the sy surface is 1250 dbar and for the s2 surface 970 dbar. A more global measure of the error between the gn surfaces and the local neutral tangent planes can be made by considering the effective diapycnal diffusivity of density when salinity and potential temperature are fluxed along gn surfaces rather than along-neutral-tangent planes. This fictitious diapycnal diffusivity 2 can be shown to be DVS ¼ Krgn z rn z , where rgn z and rnz are the two-dimensional lateral slopes of the gn surface and of the neutral tangent plane, respectively, and K is a lateral diffusivity, which we take as 1000 m2 s1. Here we have designated this false diapycnal diffusivity DVS since George Veronis and Henry Stommel first noticed this effect of mixing along a direction other than the local neutral tangent planes (they were concerned with the incorrect mixing direction being along geopotentials). In Figure 5, we have plotted the
frequency distribution of the logarithm to the base 10 of the fictitious diapycnal diffusivity DVS for various surfaces: (1) for all data in the Koltermann et al. global ocean atlas excluding the Arctic Ocean and marginal seas; and (2) for the modified Levitus reference data set. The fictitious diapycnal diffusivity was evaluated for lateral mixing along gn, sy, and s2 surfaces for (1) and only along the gn surfaces for (2) (‘labeled Levitus’ in Figure 5). The best surfaces in approximating the neutral tangent plane are clearly the two neutral density surfaces, the more accurate of these being the neutral density surface in the Levitus data that underlies the definition of neutral density. The fraction of the ocean volume exhibiting fictitious diapycnal diffusivities greater than a fixed value can be estimated by the area under these curves in Figure 5 to the right of the fixed value. It is clear that there are very relatively places where the fictitious diapycnal diffusivity exceeds 105 m2 s1 for neutral density surfaces.
Equation of State The definition of practical salinity in terms of conductivity was settled in 1978 and this is central to the evaluation of seawater density using the International Equation of State of 1980 (see Appendix 5. Properties of Seawater). These definitions and the associated algorithms have served the field of oceanography very well for almost 30 years. Recently there have been some advances in the science of determining ‘salinity’ and in evaluating many
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30
NEUTRAL SURFACES AND THE EQUATION OF STATE
0.35
0.3
0.25
2
Frequency
n 0.2
0.15
n (labeled Levitus)
0.1 n n (labeled Levitus)
0.05
0 −15 −14 −13 −12 −11 −10
−9
−8
−7
−6
−5
−4
−3
−2
−1
0
log (D) Figure 5 Frequency distribution of the logarithm of the false diapycnal diffusivity DVS arising from mixing along various coordinate surfaces.
thermodynamic properties of seawater. A working group established jointly by International Association for the Physical Sciences of the Oceans (IAPSO) and Scientific Committee on Oceanic Research (SCOR) in 2006 is examining new definitions and releases for salinity and for thermodynamic properties such as density. This is not the appropriate place to describe these developments in detail but the two recent developments that are driving this work are briefly outlined. The first development is the realization that all the accurate measurements of various thermodynamic properties of seawater (e.g., specific volume, heat capacity, sound speed) can be incorporated into a single thermodynamic function called a Gibbs function. All the thermodynamic properties of interest are obtained by simple mathematical manipulations (usually differentiation) of this Gibbs function. The salient advantages of using a Gibbs function to describe thermodynamic properties are that (1) all the derived thermodynamic properties are automatically consistent with each other and (2) the accurate data of all measurable thermodynamic properties contribute to the accuracy of many other derived thermodynamic quantities. The second development that is driving the redefinition of the thermodynamics of seawater is the realization that the differences in the composition of seawater that exist between the major ocean basins can and should be incorporated into a more accurate definition of salinity. The thermodynamic properties
of seawater depend on the ‘absolute salinity’, namely the number of grams of dissolved material in a kilogram of seawater. For seawater of standard composition, ‘absolute salinity’ is larger than ‘practical salinity’ by about 0.47%. In addition, for a given practical salinity, the absolute salinity in some ocean basins can differ by 0.02 g kg1. We expect to soon be able to estimate this spatial variation in the difference between absolute salinity and practical salinity. Since the thermodynamic properties of seawater are most consistently represented by a Gibbs function, and since this Gibbs function is best regarded as a function of absolute salinity, temperature, and pressure (rather than in terms of practical salinity), the time seems ripe to adopt a thermodynamic definition of seawater based on the Gibbs function and in terms of absolute salinity.
Summary The dynamical features in the ocean that correspond to the weather systems in the atmosphere are the socalled mesoscale eddies that typically have radii of 100 km and cause lateral mixing along neutral tangent planes with a lateral diffusivity of about 1000 m2 s1. In contrast, the diffusivity of tracers in the ocean in the diapycnal direction is 8 orders of magnitude less than this. Because of this wide disparity between the rates at which the ocean mixes
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NEUTRAL SURFACES AND THE EQUATION OF STATE
along and across ‘isopycnals’, it is important to be able to accurately determine the correct slopes of the surfaces in which the large mixing occurs so as not to have this large lateral diffusivity leak into the diapycnal direction, thus causing unwanted false diapycnal mixing. This article began by summarizing the reasons why neutral surfaces do not exist in the ocean: the neutral tangent planes that locally define the direction of neutral mixing are well defined but these planes do not link up to form a mathematically welldefined surface. Rather, oceanic motion occurs along neutral trajectories that are helical in nature and have the ability to cause a mean diapycnal advection. This diapycnal advection process is independent of the amount of diapycnal mixing present in the ocean. For the purpose of defining approximately neutral surfaces, the path-dependent nature of neutral surfaces is actually a small issue compared with the differences that exist between the hemispheres in potential density. The patched potential density procedure of Red and Lynn is relatively accurate although it is a time-consuming procedure for forming approximately neutral surfaces. The neutral density variable of Jackett and McDougall also builds in the spatial variations of water mass properties that occur in the present ocean and this software provides a practical way of forming approximately neutral surfaces in today’s ocean. The equation of state (i.e., the expression for the density of seawater in terms of the various thermodynamic parameters, commonly salinity, temperature, and pressure) was defined by UNESCO in 1981 and has served oceanography well for almost 30 years. Efforts are now underway to provide a new internationally accepted equation of state in the next year or so, this time it being based on a Gibbs potential for seawater.
See also Fluid Dynamics, Introduction, and Laboratory Experiments. Ocean Thermal Energy Conversion (OTEC). Satellite Remote Sensing of Sea Surface Temperatures. Sea Surface Exchanges of Momentum, Heat, and Fresh Water Determined by Satellite Remote Sensing. Tidal Energy. Water Types and Water Masses. Wave Energy.
31
120 g kg1. Deep Sea Research I (doi:10.1016/ j.dsr.2008.07.004). Jackett DR and McDougall TJ (1997) A neutral density variable for the world’s oceans. Journal of Physical Oceanography 27: 237--263. Koltermann KP, Gouretski V, and Jancke K (2004) Hydrographic Atlas of the World Ocean Circulation Experiment (WOCE), Vol. 3: Atlantic Ocean. http:// www.bsh.de/aktdat/mk/AIMS (accessed Mar. 2008). McDougall TJ (1987) Neutral surfaces. Journal of Physical Oceanography 17: 1950--1964. McDougall TJ (1987) Thermobaricity, cabbeling, and water-mass conversion. Journal of Geophysical Research 92: 5448--5464. McDougall TJ (1988) Neutral-surface potential vorticity. Progress in Oceanography 20: 185--221. McDougall TJ and Jackett DR (1988) On the helical nature of neutral trajectories in the ocean. Progress in Oceanography 20: 153--183. McDougall TJ and Jackett DR (2005) An assessment of orthobaric density in the global ocean. Journal of Physical Oceanography 35: 2054--2075. McDougall TJ and Jackett DR (2005) The material derivative of neutral density. Journal of Marine Research 63: 159--185. McDougall TJ and Jackett DR (2007) The thinness of the ocean in S Y p space and the implications for mean diapycnal advection. Journal of Physical Oceanography 37: 1714--1732. Millero FJ, Feistel R, Wright DG, and McDougall TJ (2008) The composition of standard seawater and the definition of the reference-composition salinity scale. Deep Sea Research I 55: 5072 (doi:10.1016/j.dsr. 2007.10.001). Phillips HB (1956) Vector Analysis, 236pp. New York: Wiley. Reid JL and Lynn RJ (1971) On the influence of the Norwegian–Greenland and Weddell seas upon the bottom waters of the Indian and Pacific oceans. Deep Sea Research and Oceanographic Abstracts 18: 1063--1088. UNESCO (1981) The Practical Salinity Scale 1978 and the International Equation of State of Seawater 1980. UNESCO Technical Papers in Marine Science 36. Wu¨st G (1933) Das Bodenwasser und die Gliedernng der Atlantischen Tiefsee. Wiss. Ergebn, Dtsch. Atlant. Exped. Meteor 6(I): 1--107
Relevant Websites http://www.cmar.csiro.au – Neutral Density, Version 3.1: Jan. 1997, CSIRO Marine and Atmospheric Research.
Further Reading Feistel R (2008) A Gibbs function for seawater thermodynamics for 6 1C to 80 1C and salinity up to
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NITROGEN CYCLE D. M. Karl, University of Hawaii at Manoa, Honolulu, HI, USA A. F. Michaels, University of Southern California, Los Angeles, CA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1876–1884, & 2001, Elsevier Ltd.
Introduction The continued production of organic matter in the sea requires the availability of the many building blocks of life, including essential major elements such as carbon (C), nitrogen (N), and phosphorus (P); essential minor elements such as iron, zinc, and cobalt; and, for many marine organisms, essential trace organic nutrients that they cannot manufacture themselves (e.g., amino acids and vitamins). These required nutrients have diverse structural and metabolic function and, by definition, marine organisms cannot survive in their absence. The marine nitrogen cycle is part of the much larger and interconnected hydrosphere–lithosphere–atmosphere–biosphere nitrogen cycle of the Earth. Furthermore, the oceanic cycles of carbon, nitrogen, and phosphorus are inextricably linked together through the production and remineralization of organic matter, especially near surface ocean phytoplankton production. This coordinated web of major bioelements can be viewed as the nutrient ‘super-cycle.’ The dominant form of nitrogen in the sea is dissolved gaseous dinitrogen (N2) which accounts for more than 95% of the total nitrogen inventory. However, the relative stability of the triple bond of N2 renders this form nearly inert. Although N2 can serve as a biologically-available nitrogen source for specialized N2 fixing microorganisms, these organisms are relatively rare in most marine ecosystems. Consequently, chemically ‘fixed’ or ‘reactive’ nitrogen compounds such as nitrate (NO3 ), nitrate (NO2 ), ammonium (NH4 þ ), and dissolved and particulate organic nitrogen (DON/ PON) serve as the principal sources of nitrogen to sustain biological processes. For more than a century, oceanographers have been concerned with the identification of growth-and production-rate limiting factors. This has stimulated investigations of the marine nitrogen cycle including both inventory determinations and pathways and controls of nitrogen transformations from one form
32
to another. Contemporaneous ocean investigations have documented an inextricable link between nitrogen and phosphorus cycles, as well as the importance of trace inorganic nutrients. It now appears almost certain that nitrogen is only one of several key elements for life in the sea, neither more nor less important than the others. Although the basic features of the marine nitrogen cycle were established nearly 50 years ago, new pathways and novel microorganisms continue to be discovered. Consequently, our conceptual view of the nitrogen cycle is a flexible framework, always poised for readjustment.
Methods and Units The analytical determinations of the various dissolved and particulate forms of nitrogen in the sea rely largely on methods that have been in routine use for several decades. Determinations of NO3 , NO2 , and NH4 þ generally employ automated shipboard, colorimetric assays, although surface waters of open ocean ecosystems demand the use of modern highsensitivity chemiluminescence and fluorometric detection systems. PON is measured by high-temperature combustion followed by chromatographic detection of the by-product (N2), usually with a commercial C–N analyzer. Total dissolved nitrogen (TDN) determination employs sample oxidation, by chemical or photolytic means, followed by measurement of NO3 . DON is calculated as the difference between TDN and the measured dissolved, reactive inorganic forms of N (NO3 , NO2 , NH4 þ ) present in the original sample. Gaseous forms of nitrogen, including N2, nitrous oxide (N2O), and nitric oxide (NO) are generally measured by gas chromatography. Nitrogen exists naturally as two stable isotopes, 14 N (99.6% by atoms) and 15N (0.4% by atoms). These isotopes can be used to study the marine nitrogen cycle by examination of natural variations in the 14N/15N ratio, or by the addition of specific tracers that are artificially enriched in 15N. Most studies of oceanic nitrogen inventories or transformations use either molar or mass units; conversion between the two is straightforward (1 mole N ¼ 14 g N, keeping in mind that the molecular weight of N2 gas is 28).
Components of the Marine Nitrogen Cycle The systematic transformation of one form of nitrogen to another is referred to as the nitrogen cycle
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NITROGEN CYCLE
(Figure 1). In the sea, the nitrogen cycle revolves around the metabolic activities of selected microorganisms and it is reasonable to refer to it as the microbial nitrogen cycle because it depends on bacteria (Table 1). During most of these nitrogen transformations there is a gain or loss of electrons and, therefore, a change in the oxidation state of nitrogen from the most oxidized form, NO3 ( þ 5), to the most reduced form, NH4 þ ( 3). Transformations in the nitrogen cycle are generally either energy-requiring (reductions) or energy-yielding (oxidations). The gaseous forms of nitrogen in the
33
surface ocean can freely exchange with the atmosphere, so there is a constant flux of nitrogen between these two pools. The natural, stepwise process for the regeneration of NO3 from PON can be reproduced in a simple ‘decomposition experiment’ in an enclosed bottle of sea water (Figure 2). During a 3-month incubation period, the nitrogen contained in particulate matter is first released as NH4þ (the process of ammonification), then transformed to NO2 (first step of nitrification), and finally, and quantitatively, to NO3 (the second step of nitrification). These transformations Atmosphere Ocean
ion
N2
fix
at
N2O
DENITRIFICATION
tro
ge
n
NH3
Ni
NO −
−
NO3 reduction
NO2 reduction
Biosynthesis
Ammonification
NO3
NO2 Ammonium oxidation
Nitrite oxidation
NITRIFICATION Export
Organic N _3
−
−
+
NH4
Import _2
_1
0 2 3 1 Oxidation state of nitrogen
4
5
Figure 1 Schematic representation of the various transformations from one form of nitrogen to another that compose the marine nitrogen cycle. Shown at the bottom is the oxidation state of nitrogen for each of the components. Most transformations are microbiological and most involve nitrogen reduction or oxidation. (Adapted from Capone, ch. 14 of Rogers and Whitman (1991).)
Table 1
Marine nitrogen cycle Credits
Process
Bacteria
Phytoplanktona
Zooplankton/Fishb
NH4þ production from DON/PON (ammonification) NH4þ /NO2 /NO3 /DON assimilation PON ingestion NH4 þ -NO2 (nitrification, step 1) NO2 -NO3 /N2O (nitrification, step 2) NO3 /NO2 -N2/N2O (denitrification) N2- NH4 þ /organic N (N2 fixation)
þ þ þ þ þ þ
þ þ þ þ
þ þ
a
Phytoplankton – eukaryotic phytoplankton. Zooplankton – including protozoans and metazoans. Abbreviations: NH4 þ , ammonium; NO2 , nitrite; NO3 , nitrate; N2O, nitrous oxide; N2, dinitrogen; DON, dissolved organic N; PON, particulate organic N; organic N, DON and PON. b
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Nitrogen (arbitrary units)
34
NITROGEN CYCLE
+
NH4
−
−
NO2
NO3
200 100
PON
0 0
1
2
and abundant planktonic cyanobacteria that coexist in tropical and subtropical marine habitats have devised alternate metabolic strategies: Synechococcus prefers NO3 and Prochlorococcus prefers NH4þ. In fact, Prochlorococcus cannot reduce NO3 to NH4þ, presumably because the critical enzyme systems are absent.
3
Time (months) Figure 2 Stepwise decomposition of particulate organic nitrogen (PON) in arbitrary units versus time during a dark incubation. Nitrogen is transformed, first to NH4þ by bacterial ammonification and finally to NO2 and NO3 by the two-step process of bacterial nitrification. These same processes are responsible for the global ocean formation of NO3 in the deep sea. These data are the idealized results of pioneering nitrogen cycle investigators, T. von Brand and N. Rakestraw, who unraveled these processes more than 50 years ago.
are almost exclusively a result of the metabolic activities of bacteria. This set of regeneration reactions is vital to the nitrogen cycle, and since most deep water nitrogen (excluding N2) is in the form of NO3, bacterial nitrification must be a very important process (see Nitrogen Distributions in the Sea, below). Nitrogen Assimilation
Several forms of nitrogen can be directly transported across cell membranes and assimilated into new cellular materials as required for biosynthesis and growth. Most microorganisms readily transport NH4þ, NO2, NO3, and selected DON compounds such as amino acids, urea, and nucleic acid bases. By comparison, the ability to utilize N2 as a nitrogen source for biosynthesis is restricted to a very few species of specialized microbes. Many protozoans, including both photosynthetic and heterotrophic species, and all metazoans obtain nitrogen primarily by ingestion of PON. Once inside the cell or organism, nitrogen is digested and, if necessary, reduced to NH4þ. If oxidized compounds such as NO3 or NO2 are utilized, cellular energy must be invested to reduce these substrates to ammonium for incorporation into organic matter. The process of reduction of NO3 (or NO2) for the purpose of cell growth is referred to as assimilatory nitrogen (NO3/NO2) reduction and most microorganisms, both bacteria and phytoplankton, possess this metabolic capability (Table 1). In theory, there should be a metabolic preference for NH4þ over either NO3 or NO2, based strictly on energetic considerations. However, it should be emphasized that preferential utilization of NH4þ does not always occur. For example, two closely related
Nitrification
As nitrogen is oxidized from NH4þ through NO2 to NO3, energy is released (Figure 1), a portion of which can be coupled to the reduction of carbon dioxide (CO2) to organic matter (CH2O) by nitrifying bacteria. These specialized bacteria, one group capable only of the oxidation of NH4þ to NO2 and the second capable only of the oxidation of NO2 to NO3, are termed ‘chemolithoautotrophic’ because they can fix CO2 in the dark at the expense of chemical energy. Other related chemolithoautotrophs can oxidize reduced sulfur compounds, and this pathway of organic matter production has been hypothesized as the basis for life at deep-sea hydrothermal vents. It is essential to emphasize an important ecological aspect of NH4þ/NO2 chemolithoautotrophy. First, the oxidation of NH4þ to NO2 and of NO2 to NO3 usually requires oxygen and these processes are ultimately coupled to the photosynthetic production of oxygen in the surface water. Second, the continued formation of reduced nitrogen, in the form of NH4þ or organic nitrogen, is also dependent, ultimately, on photosynthesis. In this regard the CO2 reduced via this ‘autotrophic’ pathway must be considered secondary, not primary, production from an ecological energetics perspective. Marine nitrifying bacteria, especially the NO2 oxidizers are ubiquitous in the world ocean and key to the regeneration of NO3, which dominates waters below the well-illuminated, euphotic zone. However they are never very abundant and, at least for those species in culture, grow very slowly. Certain heterotrophic bacteria can also oxidize NH4þ to both NO2 and NO3 during metabolism of preformed organic matter. However, very little is known about the potential for ‘heterotrophic nitrification’ in the sea. Denitrification
Under conditions of reduced oxygen (O2) availability, selected species of marine bacteria can use NO3 as a terminal acceptor for electrons during metabolism, a process termed NO3 respiration or dissimilatory NO3 reduction. This process allows microorganisms to utilize organic matter in low-O2
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NITROGEN CYCLE
or anoxic habitats with only a slight loss of efficiency relative to O2-based metabolism. A majority of marine bacteria have the ability for NO3 respiration under the appropriate environmental conditions (Table 1). Potential by-products of NO 3 respiration are NO2, N2, and N2O; if a gas is formed (N2/N2O) then the process is termed denitrification because the net effect is to remove bioavailable nitrogen from the local environment. The total rate of denitrification is generally limited by the availability of NO3, and a continued supply of NO3 via nitrification is dependent upon the availability of NH4þ and free O2. Consequently denitrification typically occurs at boundaries between low-O2 and anoxic conditions where the supply of NH4þ from the anoxic zone sustains a high rate of NO3 production via nitrification to fuel- sustained NO3 respiration and denitrification. Recently a new group of microorganisms has been isolated that are capable of simultaneously using both O2 and NO3/NO2 as terminal electron acceptors. This process is termed ‘aerobic denitrification.’ Likewise, there are exceptional microorganisms that are able to carry out anaerobic nitrification (oxidation of NH4þ in the absence of O2). It appears difficult to establish any hard-and-fast rules regarding marine nitrogen cycle processes. N2 Fixation
The ability to use N2 as a growth substrate is restricted to a relatively small group of microorganisms. Open ocean ecosystems that are chronically depleted in fixed nitrogen would appear to be ideal habitats for the proliferation of N2-fixing microorganisms. However, the enzyme that is required for reduction of N2 to NHþ 4 is also inhibited by O2, so specialized structural, molecular, and behavioral adaptations have evolved to promote oceanic N2 fixation. Fixation of molecular nitrogen in the open ocean may also be limited by the availability of iron, which is an essential cofactor for the N2 reduction enzyme system. Changes in iron loading are caused by climate variations, in particular the areal extent of global deserts, by the intensity of atmospheric circulation, and more recently by changes in land use practices. Conversion of deserts into irrigated croplands may cause a change in the pattern and intensity of dust production and, therefore, of iron transport to the sea. Humanity is also altering the global nitrogen cycle by enhancing the fixation of N2 by the manufacturing of fertilizer. At the present time, the industrial fixation of N2 is approximately equivalent to the pre-industrial, natural N2 fixation rate.
35
Eventually some of this artificially fixed N2 will make its way to the sea, and this may lead to a perturbation in the natural nitrogen cycle. On a global scale and over relatively long timescales, the total rate of N2 fixation is more or less in balance with total denitrification, so that the nitrogen cycle is mass-balanced. However, significant net deficits or excesses can be observed locally or even on ocean basin space scales and on decade to century timescales. These nitrogen imbalances may impact the global carbon and phosphorus cycles as well, including the net balance of CO2 between the ocean and the atmosphere.
Nitrogen Distributions in the Sea Required growth nutrients, like nitrogen, typically have uneven distributions in the open sea, with deficits in areas where net organic matter is produced and exported, and excesses in areas where organic matter is decomposed. For example, surface ocean NO3 distributions in the Pacific basin reveal a coherent pattern with excess NO3 in high latitudes, especially in the Southern Ocean (south of 601 S), and along the Equator (especially east of the dateline), and generally depleted NO3 concentrations in the middle latitudes of both hemispheres (Figure 3). These distributions are a result of the balance between NO3 supply mostly by ocean mixing and NO3 demand or net photosynthesis. The very large NO3 inventory in the surface waters of the Southern Ocean implies that factors other than fixed nitrogen availability control photosynthesis in these regions. It has been hypothesized that the availability of iron is key in this and perhaps other regions of the open ocean. The much smaller but very distinctive band of elevated NO3 along the Equator is the result of upwelling of NO3-enriched waters from depth to the surface. This process has a large seasonal and, especially, interannual variability, and it is almost absent during El Nin˜o conditions. Excluding these high-latitude and equatorial regions, the remainder of the surface waters of the North and South Pacific Oceans from about 401N to 401S are relatively depleted in NO3. In fact surface (0–50 m) NO3 concentrations in the North Pacific subtropical gyre near Hawaii are typically below 0.01 mmol l1 (Figure 4). Within the upper 200 m, the major pools of fixed nitrogen (e.g., NO3, DON, and PON) have different depth distributions. In the sunlit surface zone, NO3 is removed to sustain organic matter production and export. Beneath 100 m, there is a steep concentration versus depth gradient (referred to as the nutricline), which reaches a
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36
NITROGEN CYCLE
70˚N 60˚N 50˚N 40˚N 30˚N 20˚N
Latitude
10˚N 0˚ 10˚S 20˚S 30˚S 40˚S 50˚S 60˚S 70˚S 80˚S 105˚E 120˚E 135˚E 150˚E
165˚E 180˚E 165˚W 150˚W 135˚W 120˚W 105˚W Longitude
90˚W
75˚W
60˚W
Figure 3 Mean annual NO3 concentration (mmol l1) at the sea surface for samples collected in the Pacific Ocean basin and Pacific sector of the Southern Ocean. (From Conkright et al. 1998.)
maximum of about 40–45 it mmol l1 at about 1000 m in the North Pacific Ocean. PON concentration is greatest in the near-surface waters where the production of organic matter via photosynthesis is highest (Figure 4). PON includes both living (biomass) and nonliving (detrital) components; usually biomass nitrogen is less than 50% of the total PON in near-surface waters, and less than 10% beneath the euphotic zone (4150 m). DON concentration is also highest in the euphotic zone (B5– 6 mmol l1) and decreases systematically with depth to a minimum of 2–3 mmol l1 at 800–1000 m. The main sources for DON in the surface ocean are the combined processes of excretion, grazing, death, and cell lysis. Consequently, DON is a complex mixture of cell-derived biochemicals; at present, less than 20% of the total DON has been chemically characterized. Dissolved N2 (not shown) is always high (B800 mmol l1) and increases systematically with depth. The major controls of N2 concentration are temperature and salinity, which together determine gas solubility. Marine life has little impact on N2 distributions in the open sea even though some microorganisms can utilize N2 as a growth substrate and others can produce N2 as a metabolic by-
product. These transformations are simply too small to significantly impact the large N2 inventories in most regions of the world ocean. Another important feature of the global distribution of NO3 is the regional variability in the deep water inventory (Figure 5). Deep ocean circulation can be viewed as a conveyor-belt-like flow, with the youngest waters in the North Atlantic and the oldest in the North Pacific. The transit time is in excess of 1000 y, during which time NO3 is continuously regenerated from exported particulate and dissolved organic matter via coupled ammonification and nitrification (Figure 2). Consequently, the deep Pacific Ocean has nearly twice as much NO3 as comparable depths in the North Atlantic (Figure 5). Nitrous Oxide Production
Nitrous oxide (N2O) is a potent greenhouse gas that has also been implicated in stratospheric ozone depletion. The atmospheric inventory of N2O is presently increasing, so there is a renewed interest in the marine ecosystem as a potential source of N2O. Nitrous oxide is a trace gas in sea water, with typical concentrations ranging from 5 to 50 nmol l1.
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NITROGEN CYCLE
_1
_1
Nitrate (μmol l )
37
_1
DON (μmol l )
PON (μmol l )
0 20 40 60
Depth (m)
80 100 120 140 160 180 200 0
0.5
1.0
1.5 4.5
5.0
6.0 0
5.5
0.1
0.2
0.3
0.4
Figure 4 Average concentrations (mmol l1) of NO3, DON, and PON versus water depth for samples collected in the upper 200 m of the water column at Sta. ALOHA (22.751N, 158.01W). These field data are from the Hawaii Ocean Time-series program and are available at http://hahana.soest.edu/hot_jgofs.html).
0
Primary Nitrite (NO 2 ) Maximum
An interesting, almost cosmopolitan feature of the world ocean is the existence of a primary NO2
1000
N. Pacific
N. Atlantic 2000
Depth (m)
Concentrations of N2O in oceanic surface waters are generally in slight excess of air saturation, implying both a local source and a sustained ocean-to-atmosphere flux. Typically there is a mid-water (500– 1000 m) peak in N2O concentration that coincides with the dissolved oxygen minimum. At these intermediate water depths, N2O can exceed 300% saturation relative to atmospheric equilibrium. The two most probable sources of N2O in the ocean are bacterial nitrification and bacterial denitrification, although to date it has been difficult to quantify the relative contribution of each pathway for a given habitat. Isotopic measurements of nitrogen and oxygen could prove invaluable in this regard. Because the various nitrogen cycle reactions are interconnected, changes in the rate of any one process will likely have an impact on the others. For example, selection for N2-fixing organisms as a consequence of dust deposition or deliberate iron fertilization would increase the local NHþ 4 inventory and lead to accelerated rates of nitrification and, hence, enhanced N2O production in the surface ocean and flux to the atmosphere.
3000
4000
5000
6000 0
5
10
15
20
25
30
35
40
45
_1
Nitrate (μmol l ) Figure 5 Nitrate concentrations (mmol l1) versus water depth at two contrasting stations located in the North Atlantic (31.81N, 50.91W) and North Pacific (301N, 160.31W) Oceans. These data were collected in the 1970s during the worldwide GEOSECS expedition, stations #119 and #212, respectively.
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38
NITROGEN CYCLE
maximum (PNM) near the base of the euphotic zone (B100–150 m; Figure 6). Nitrite is a key intermediate between NO3 and NH4þ, so there are several potential pathways, both oxidative and reductive, that might lead to its accumulation in sea water. First, phototrophic organisms growing on NO3 may partially reduce the substrate to NO2 as the first, and least energy-consuming, step in the assimilatory NO 3 reduction pathway. However, the next step, reduction of NO2 to NH4þ, requires a substantial amount of energy, so when energy is scarce (e.g., light limitation) NO2 accumulates inside the cells. Because NO2 is the salt of a weak acid, nitrous acid (HNO2) forms in the slightly acidic intracellular environment, diffuses out of the cell and ionizes to form NO2 in the alkaline sea water. This NO3-NO2 phytoplankton pump, under the control of light intensity, could provide a source of NO 2 necessary to create and maintain the PNM. Alternatively, local regeneration of dissolved and particulate organic matter could produce NH4þ (via ammonification) that is partially oxidized in place to produce a relative excess of NO2 (the first step of
nitrification). Kinetic controls on this process would be rates of NH4þ production and NO2 oxidation to NO3 (the second and final step in nitrification). Sunlight, even at very low levels, appears to disrupt the normal coupling between NO2 production and NO2 oxidation, in favor of NO2 accumulation. Finally, it is possible, though perhaps less likely, that NO3 respiration (terminating at NO2), followed by excretion of NO2 (into the surrounding sea water might also contribute to the accumulation of NO2) near the base of the euphotic zone. Because the global ocean at the depth of the PNM is characteristically well-oxygenated, one would need to invoke microenvironments like animal guts or large particles as the habitats for this nitrogen cycle pathway. The use of 15N-labeled substrates, selective metabolic inhibitors, and other experimental manipulations provides an opportunity for direct assessment of the role of each of these potential processes. In all likelihood, more than one of these processes contributes to the observed PNM. Whatever the cause, light appears to be an important determinant that might explain the relative position, with regard to depth, of this global feature.
0
Nitrogen Cycle and Ocean Productivity
Light
40
Depth (m)
80 Chlorophyll-a
120 −
− NO2
NO3
160
200 Concentration Figure 6 Schematic representation of the depth distributions of sunlight, chlorophyll-a, NO 2 , and NO3 for a representative station in the subtropical North Pacific Ocean showing the relationship of the primary NO 2 maximum (PNM) zone (shaded) to the other environmental variables. (Modified from J. E. Dore, Microbial nitrification in the marine euphotic zone, Ph.D. Dissertation, University of Hawaii, redrawn with permission of the author.)
Because nitrogen transformations include both the formation and decomposition of organic matter, much of the nitrogen used in photosynthesis is locally recycled back to NH4þ or NO3 to support another pass through the cycle. The net removal of nitrogen in particulate, dissolved, or gaseous form can cause the cycle to slow down or even terminate unless new nitrogen is imported from an external source. A unifying concept in the study of nutrient dynamics in the sea is the ‘new’ versus ‘regenerated’ nitrogen dichotomy (Figure 7). New nitrogen is imported from surrounding regions (e.g., NO3 injection from below) or locally created (e.g., NH4þ/organic N from N2 fixation). Regenerated nitrogen is locally recycled (e.g., NH4þ from ammonification, NO2 / NO3 from nitrification, or DON from grazing or cell lysis). Under steady-state conditions, the amount of new nitrogen entering an ecosystem will determine the total amount that can be exported without the system running down. In shallow, coastal regions runoff from land or movement upward from the sediments are potentially major sources of NH4þ, NO3 and DON for water column processes. In certain regions, atmospheric deposition (both wet and dry) may also supply bioavailable nitrogen to the system. However, in
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NITROGEN CYCLE
39
N2
Gas exchange
N2
Atmosphere
Euphotic zone
N2 fixation PRIMARY PRODUCERS
CONSUMERS
−
NO3
+
NH4
Recycling DON Nitracline Export
Export Ammonification
Diffusion and advection
DON and PON
+
NH4
DON and PON Aphotic zone
−
NO3
Nitrification Figure 7 Schematic representation of the major pools and transformations/fluxes of nitrogen in a typical open ocean ecosystem. New sources of bioavailable N (NO3 and N2 in this presentation) continuously resupply nitrogen that is lost via DON and PON export. These interactions and ocean processes form the conceptual framework for the ‘new’ versus ‘regenerated’ paradigm of nitrogen dynamics in the sea that was originally proposed by R. Dugdale and J. Goering.
most open ocean environments, new sources of nitrogen required to balance the net losses from the euphotic zone are restricted to upward diffusion or mixing of NO3 from deep water and to local fixation of N2 gas. In a balanced steady state, the importation rate of new sources of bioavailable nitrogen will constrain the export of nitrogen (including fisheries production and harvesting). If all other required nutrients are available, export-rich ecosystems are those characterized by high bioavailable nitrogen loading such as coastal and open ocean upwelling regions. These are also the major regions of fish production in the sea.
See also Atmospheric Input of Pollutants. Nitrogen Isotopes in the Ocean. Phosphorus Cycle. Primary Production Processes.
Further Reading Carpenter EJ and Capone DG (eds.) (1983) Nitrogen in the Marine Environment. New York: Academic Press. Conkright ME, O’Brien TD, Levitus S, et al. (1998) NOAA Atlas NESDIS 37, WORLD OCEAN ATLAS 1998, vol. II: Nutrients and Chlorophyll of the Pacific Ocean. Washington, DC: US Department of Commerce. Harvey HW (1966) The Chemistry and Fertility of Sea Waters. London: Cambridge University Press. Kirchman DL (ed.) (2000) Microbial Ecology of the Oceans. New York: Wiley-Liss. Rogers JE and Whitman WB (eds.) (1991) Microbial Production and Consumption of Greenhouse Gases: Methane, Nitrogen Oxides, and Halomethanes. Washington, DC: American Society for Microbiology. Schlesinger WH (1997) Biogeochemistry: An Analysis of Global Change. San Diego: Academic Press. Wada E and Hattori A (1991) Nitrogen in the Sea: Forms, Abundances, and Rate Processes. Boca Raton, FL: CRC Press.
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NITROGEN ISOTOPES IN THE OCEAN D. M. Sigman and K. L. Karsh, Princeton University, Princeton, NJ, USA K. L. Casciotti, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Nitrogen has two stable isotopes, 14N and 15N (atomic masses of 14 and 15, respectively). 14N is the more abundant of the two, comprising 99.63% of the nitrogen found in nature. Physical, chemical, and biological processes discriminate between the two isotopes, leading to subtle but measurable differences in the ratio of 15N to 14N among different forms of nitrogen found in the marine environment. Nitrogen is a central component of marine biomass and one of the major nutrients required by all phytoplankton. In this sense, biologically available (or ‘fixed’, i.e., non-N2) N is representative of the fundamental patterns of biogeochemical cycling in the ocean. However, N differs from other nutrients in that its oceanic sources and sinks are dominantly internal and biological, with marine N2 fixation supplying much of the fixed N in the ocean and marine denitrification removing it. The N isotopes provide a means of studying both the input/output budget of oceanic fixed N and its cycling within the ocean. In this overview, we outline the isotope systematics of N cycle processes and their impacts on the isotopic composition of the major N reservoirs in the ocean. This information provides a starting point for considering the wide range of questions in ocean sciences to which the N isotopes can be applied.
Terms and Units Mass spectrometry can measure precisely the ratio of the N isotopes relative to a N reference containing a constant isotopic ratio. The universal reference for N isotopes is atmospheric N2, with an 15N/14N ratio of 0.367 65%70.000 81%. Natural samples exhibit small deviations from the standard ratio, which are expressed in d notation (in units of per mil, %):
15
d Nð%Þ ¼
40
ð15 N=14 NÞsample ð15 N=14 NÞstandard
! 1
1000 ½1
In this notation, the d15N of atmospheric N2 is 0%. Special terms are also used to characterize the amplitude of isotopic fractionation caused by a given process. Isotope fractionation results from both equilibrium processes (‘equilibrium fractionation’) and unidirectional reactions (‘kinetic fractionation’). Nitrogen isotope variations in the ocean are typically dominated by kinetic fractionation associated with the conversions of N from one form to another. The kinetic isotope effect, e, of a given reaction is defined by the ratio of rates with which the two N isotopes are converted from reactant to product: eð%Þ ¼ ð14 k=15 k 1Þ 1000
½2
where 14k and 15k are the rate coefficients of the reaction for 14N- and 15N-containing reactant, respectively. For e{1000%, e is approximated by the difference in d15N between the reactant and its instantaneous product. That is, if a reaction has an e of 5%, then the d15N of the product N generated at any given time will be B5% lower than the d15N of the reactant N at that time.
Measurements The isotopic analysis of N relies on the generation of a stable gas as the analyte for isotope ratio mass spectrometry. Online combustion to N2 is currently the standard method for the preparation of a N sample for isotopic analysis. With ‘off-the-shelf’ technology, a typical sample size requirement is 1–2 mmol N per analysis. Gas chromatography followed by combustion to N2 is improving as a technique for specific organic compounds, amino acids in particular, although the polarity of many N compounds remains a challenge. Liquid chromatography is also being explored. There are standard methods of collection for most bulk forms of particulate N (PN) in the ocean. Shallow and deep samples of suspended PN are filtered onto glass fiber filters. Sinking PN is collected by sediment traps. Zooplankton can be picked from filtered samples or net tows, and particulates can be separated into size classes. In the case of dissolved forms of N, the species of interest must be converted selectively to a gas or other extractable form for collection. Since the 1970s, the d15N values of marine nitrate (NO3 ), nitrite (NO2 ), and ammonium (NH4 þ ) have been analyzed by conversion to ammonia gas and collection of the cationic ammonium form for
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41
NITROGEN ISOTOPES IN THE OCEAN
Models Two simple models, the ‘Rayleigh’ model and the ‘steady-state’ model, are frequently used to interpret N isotope data from the ocean. In both of these models, the degree of consumption of the reactant N pool (f ) is a key parameter, and the d15N of the initial reactant N pool (d15 Ninitial ) and kinetic isotope effect (e) are the two central isotopic parameters. If a transformation proceeds with a constant isotope effect and if the reactant and product N pools are neither replenished nor lost from the system during the progress of the transformation, then the process can be described in terms of Rayleigh fractionation kinetics, which define the isotopic variation of the reactant N pool (d15 Nreactant ; eqn [3]), the instantaneously generated product N (d15 Ninstantaneous ; eqn [4]), and the integrated product N pool (d15 Nintegrated ; eqn [5]) as a given reservoir of reactant N is consumed (Figure 1): 15
15
d Nreactant ¼ d Ninitial eflnðf Þg 15
15
d Ninstantaneous ¼ d Nreactant e
½3
25
Rayleigh (closed system): 15Nreactant 15Ninstantaneous
20 15N (% vs. air)
subsequent conversion to N2 (often referred to as the ammonia ‘distillation’ and ‘diffusion’ methods). Recently, more sensitive isotope analysis methods (requiring only 5–10 nmol of N per analysis) have been developed for nitrate and nitrite in which these species are converted to nitrous oxide (N2O), followed by isotopic analysis of this gas (the ‘bacterial’ or ‘denitrifier’ method and the ‘chemical’ or ‘azide’ method). The N2O produced by these methods (or naturally occuring N2O) is analyzed by a purge and trap system, followed by gas chromatography and isotope ratio mass spectrometry. The N2O-based methods also allow for oxygen isotope analysis of nitrate and nitrite, a measurement not previously possible in seawater. In addition, they provide a cornerstone for isotopic analysis of other dissolved forms of N, such as dissolved organic N (DON) and ammonium (NH4 þ ), which can be converted to nitrate and/or nitrite. With respect to dissolved gases, methods of collection and isotopic analysis have been developed for N2 and N2O, with recent progress on isotopomer analysis of N2O (i.e., distinguishing 15N14N16O from 14N15N16O).
15Nintegrated 15
Steady-state (open system): 15Nreactant 15Nproduct
10 5 0 0
0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 Fraction of reactant remaining
1
Figure 1 The d15N of reactant and product N pools of a single unidirectional reaction as a function of the fraction of the initial reactant supply that is left unconsumed, for two different models of reactant supply and consumption, following the approximate equations given in the text. The Rayleigh model (black lines) applies when a closed pool of reactant N is consumed. The steady-state model (gray lines) applies when reactant N is supplied continuously. The same isotopic parameters, an isotope effect (e) of 5% and a d15N of 5% for the initial reactant supply, are used for both the Rayleigh and steady-state models. e is approximately equal to the isotopic difference between reactant N and its product (the instantaneous product in the case of the Rayleigh model).
forms of the full expressions. They are typically adequate, but their error is greater for higher consumption (lower f ) and higher e. The Rayleigh model is often used to describe events in the ocean, such as the uptake of nitrate by phytoplankton during a bloom in a stratified surface layer. The end-member alternative to the Rayleigh model is the steady-state model, in which reactant N is continuously supplied and partially consumed, with residual reactant N being exported at a steady-state rate such that the gross supply of reactant N equals the sum of the product N and the residual reactant N exported. In this case, the following approximate expessions apply to the reactant N pool (d15 Nreactant ; eqn [6]) and the product N pool (d15 Nproduct ; eqn [7]) (Figure 1): d15 Nreactant ¼ d15 Ninitial þ eð1 f Þ
½6
d15 Nproduct ¼ d15 Ninitial eðf Þ
½7
½4
d15 Nintegrated ¼ d15 Ninitial þ eff =ð1 f Þglnðf Þ ½5 where f is the fraction of the reactant remaining, d15Ninitial is the d15N of the initial reactant N pool, and e is the kinetic isotope effect of the transformation. These equations are simplified, approximate
The steady-state model and modified forms of it, such as the more spatially complex ‘reaction– diffusion’ model, are used to quantify uptake processes where supply and uptake are simultaneous and relatively time-invariant, such as in the consumption of nitrate by denitrification in the ocean interior or in sediments.
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42
NITROGEN ISOTOPES IN THE OCEAN
Processes Inputs
N2 fixation is the major input of fixed N to the ocean (Figure 2). N2 fixation is carried out by N2 fixers, cyanobacteria and other microorganisms able to convert N2 into biomass N. Subsequent remineralization of this biomass supplies new N to the dissolved fixed N pools in the surface and subsurface ocean. Field collections of Trichodesmium colonies, the best-known genus of open ocean N2 fixer, have yielded a d15N of c. 2% to þ 0.5%. Taking into account the d15N of dissolved N2 (0.6% in the surface mixed layer), this range in d15N is consistent with, but perhaps less variable than, the range in isotope effects estimated from culture studies of marine and terrestrial N2 fixers, B0–4% (Table 1). An average d15N of 1% has been suggested for the fixed N input to the ocean from N2 fixation.
Other inputs of fixed N to the marine environment include terrestrial runoff and atmospheric precipitation, the N isotopic compositions of which are poorly constrained (Figure 2). Dissolved and particulate d15N in pristine river systems ranges mostly from 0% to 5%. However, biological processes along the flow path and in estuaries (in particular, by denitrification; see below) can alter the d15N of the final inputs from terrestrial runoff in complex ways. Anthropogenic inputs often increase the d15N of a system because they encourage denitrification. In atmospheric inputs, a wide range in the d15N of inorganic (c. 16% to 10%) and organic (c. 8% to 1%) N has been observed, with increasing evidence that at least some of this variability can provide insight into sources and processes. In the face of large uncertainties, a preindustrial mean d15N of 4% for terrestrial runoff and 2% for atmospheric precipitation has been suggested by some workers.
Sinking 15N ~ 3.5‰
N2 fixation
≤ 2‰
Atmospheric precipitation 15N ~ −2‰?
Terrestrial runoff 15N ~ 4‰?
Dissolved N2 15N = 0.6‰
~ 5‰
− supply and as si m ila t
NO 3
Thermocline NO3− 15N ~ 3‰
Particulate N 15N ~ 0‰
≤ 3‰?
≤ 3‰? DON 15N ~ 4‰
?
ion
Deep ocean
Recycling
NH4+
Remineralization
Surface ocean
Atmospheric N2 15N = 0‰
Water-column denitrification ~ 25‰
Sedimentary denitrification ~ 0‰
Deep NO3− 15N = 5‰
Figure 2 Processes affecting the distribution of nitrogen isotopes in the sea. The inputs and outputs (solid arrows) control the ocean’s inventory of fixed N, the majority of which is in the form of nitrate (NO3 ). Inputs are marine N2 fixation in the surface ocean, terrestrial runoff, and atmospheric precipitation. Outputs indicated are sedimentary and water column denitrification. As discussed in the text, water-column denitrification in low-oxygen regions of the ocean interior leads to elevated d15N of nitrate (dark area). Remineralization of newly fixed N can explain the low d15N of nitrate in the shallow subsurface, or thermocline, of low-latitude regions (light area). Internal cycling is represented with dashed arrows. Nitrate supplied from the deep ocean and thermocline is assimilated in the surface ocean. PN is recycled in the surface ocean, degraded to ammonium (NH4 þ ) that is subsequently assimilated; the role of DON in this recycling is not yet clear. Sinking and remineralization (via degradation and nitrification, see text) returns N from the particulate pool to nitrate. For simplicity, possible nitrification in the surface ocean is not shown. The isotope fractionation associated with nitrification in the ocean interior is also excluded because this process generally goes to completion (see text). The d15N of particulate suspended and sinking N, DON, and thermocline nitrate is taken from the subtropical North Atlantic. Question marks indicate the greatest uncertainties, due to variation in the available data and/or insufficient data.
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NITROGEN ISOTOPES IN THE OCEAN
Table 1
43
Representative estimates of isotope effects for N cycles processes
Process
Isotope effect (e)
Details
N2 fixation (N2-PN)
1.8–3.0% 0.2% 0.4–0.2% 5–30% 18–28% 25 30% r3% 5–17% 5–6% 5–9% 5% 5% 20% 8–27% 4–27% 6.5–8% 9% 18.5%
Trichodesmium spp. Western Tropical North Pacific Western Tropical Atlantic Pseudomonas stutzeri (marine) Paracoccus denitrificans (terrestrial) Eastern Tropical North Pacific and Arabian Sea Coastal sites, eastern N. Pacific margin, and deep Bering Sea Thalassiosira weissflogii Coastal Antarctic Open Antarctic and Subantarctic Subarctic Pacific Equatorial Pacific Thalassiosira pseudonana Skeletonema costatum Bacterial assemblage Chesapeake Bay Delaware Estuary Scheldt Esturary
14% 19% 35–38% 12–16% Unknown
Nitrosomonas marina (marine) Nitrosomonas C-113a (marine) Nitrosomonas europaea (terrestrial) Chesapeake Bay
Denitrification (NO3 -N2) Water column Sedimentary NO3 assimilation (NO3 -PN)
NH4 þ assimilation (NH4 þ -PN)
Nitrification Ammonia oxidation (NH3 -NO2 )
Nitrite oxidation (NO2 -NO3 )
Outputs
Denitrification, the bacterial reduction of nitrate to N2, is the major mechanism of fixed N loss from the ocean, occurring both in the water column and in sediments when the oxygen concentration is low (o5 mM) (Figure 2). Denitrification strongly discriminates against the heavier isotope, 15N, progressively enriching the remaining nitrate pool in 15N as nitrate consumption proceeds. Culture studies of denitrifying bacteria suggest (with some exceptions) an isotope effect of B20–30%, a range supported by water column estimates (Table 1). The isotopic discrimination during denitrification likely takes place as nitrate is reduced intracellularly to nitrite by the dissimilatory form of the enzyme nitrate reductase, such that unconsumed, 15N-enriched nitrate effluxing from the cell back into ambient waters allows the enzyme-level isotope effect to be expressed at the environmental scale. Where it occurs in low-oxygen regions of the mid-depth ocean, water column denitrification causes a clear elevation in the d15N of nitrate, and it is the reason that global ocean nitrate d15N is higher than that of the N source from N2 fixation, the dominant input. In contrast to water-column denitrification, denitrification in sediments leads to little increase in the d15N of water-column nitrate. The high d15N of nitrate within the pore waters of actively denitrifying sediments demonstrates that isotopic discrimination
occurs at the scale of the organism. However, expression of the organism-scale isotope effect at the scale of sediment/water exchange is minimized by nearly complete consumption of the nitrate at the site of denitrification within sediment pore waters, which prevents 15N-enriched residual nitrate from evading to the overlying water column, yielding an ‘effective’ isotope effect of 3% or less in most sedimentary environments studied so far (Table 1). Another mechanism of fixed N loss that occurs in sediments and the water column is anaerobic ammonium oxidation, or ‘anammox’, in which nitrite (from nitrate reduction or ammonium oxidation) is used to oxidize ammonium to N2 (NO2 þ NH4 þ -N2 þ 2H2 O). This process has unknown effects on isotope distributions in the ocean. The effects of anammox on N isotopes must depend on the organism-scale isotope effects, the sources of nitrite and ammonium substrates for the reaction, and the degree to which these substrates are consumed. For instance, if nitrate reduction by denitrifiers is the source of the nitrite, remineralization processes are the source of the ammonium, and both the nitrite and ammonium are completely consumed in the environment where anammox occurs, then the isotope discrimination would simplify to that of the nitrate reduction by denitrifiers averaged with any isotope discrimination during the remineralization that produces the needed ammonium. It should be noted that many water-column-derived
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44
NITROGEN ISOTOPES IN THE OCEAN
isotope effect estimates for ‘denitrification’ have inherently included the effect of anammox, in that they regress the nitrate d15N increase against the total nitrate deficit relative to phosphate, and ammonium is not observed to accumulate. Internal Cycling
The fluxes associated with internal cycling are neither sources nor sinks of fixed N but affect the distributions of N species and isotopes in the ocean. N assimilation In the surface ocean, phytoplankton assimilate fixed N (nitrate and ammonium, as well as nitrite, urea, and other organic N compounds) (Figure 3). Culture studies indicate that different forms of fixed N are assimilated with distinct isotope effects, although these isotope effects may vary with physiological conditions. For all studied forms, phytoplankton preferentially consume 14N relative to 15 N (Figures 3 and 4). Nitrate is the deep-water source of fixed N for phytoplankton growth, and the degree of its consumption varies across the surface ocean. The isotope effect of nitrate assimilation therefore has a major impact on the isotopic distributions of all N forms in the ocean. Field-based estimates of the isotope effect of nitrate assimilation range from 4% to 10%, with most estimates closer to 5–8% (Table 1).
PN
Culture-based estimates are more variable. Physiological studies suggest that isotopic fractionation associated with nitrate assimilation is imparted by the intracellular assimilatory nitrate reductase enzyme, which has an estimated intrinsic isotope effect of 15– 30%. The enzyme-level isotope effect is expressed by efflux of unconsumed nitrate out of the cell, as also appears to be the case with denitrifiers. The lower range of isotope effect estimates associated with algal nitrate assimilation than with denitrification suggests proportionally less nitrate efflux by algae, perhaps related to the fact that fixed N is often scarce in the surface ocean. Some studies suggest the degree of efflux and isotope effect of nitrate assimilation may vary with growth conditions; one set of studies of the diatom Thalassiosira weissflogii showed a higher isotope effect under light-limited growth than under growth limited by iron or temperature. The isotope effect of nitrate assimilation is an integrative characteristic of the upper ocean biota that can be measured without perturbing the system, and its magnitude is of broad significance in the application of N isotopes to various questions in the modern and past oceans. Thus, it is a priority to develop a predictive understanding of its controls. Other forms of fixed N assimilated by phytoplankton (ammonium, nitrite, and urea) are
Degradation ( ≤ 3‰ ?) DON
Assimilation (nitrate ~ 5‰)
NH4+
(nitrite ~ 1‰)
(ammonium 20‰) (urea ~ 1‰) NO2−
Nitrification (ammonia oxidation ~ 16‰) (nitrite oxidation ~ ?)
NO3−
Figure 3 Schematic diagram of the processes and pools central to the internal cycling of N in the ocean. The isotope effects shown here are based on laboratory studies. Dashed arrows represent assimilation of dissolved species into particulate matter, and solid arrows represent remineralization. Complete consumption of the ammonium pool by assimilation in the surface ocean or by nitrification in the ocean interior causes the relatively high isotope effects associated with these processes to have little effect on N isotope dynamics. However, in regions where ammonium assimilation and nitrification co-occur, their isotope effects will impact the d15N of their respective products, PN and nitrate. In nitrification, ammonia (NH3), rather than the protonated form ammonium (NH4þ ), is oxidized. However, ammonium is the dominant species in seawater, and there is isotope discrimination in the ammonium–ammonia interconversion. Thus, the isotope effects for ‘ammonia oxidation’ given here and elsewhere in the text refer specifically to consumption of ammonium. The processes surrounding DON production and utilization are not well understood from an isotopic perspective but are thought to play an important role in N cycling.
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NITROGEN ISOTOPES IN THE OCEAN
30 NO3−
150
25 Fluor.
20 15
100
10 50
5
15N-PN(‰)
4
5
4
5 0
0 (b)
(c) 6 Fluorescence
PN
15N-PN(‰)
Nitrogen (μM)
(a) 200
45
3
2
3 15N of source
2
1 15N = 4.4−4.5 F
1 0 0
30
40
50 Time (h)
60
70
0 1.0
0.5 F
0
Figure 4 Isotopic fractionation during nitrate (NO3 ) assimilation by a culture of the marine diatom Thalassiosira pseudonana. (a) Time series of NO3 concentration in the culture medium (filled circles), PN (open circles), and fluorescence, a measure of phytoplankton biomass (filled triangles); (b) d15N of PN accumulated over the same time period; (c) d15N of PN accumulated during log phase of growth plotted vs. F, a measure of nitrate utilization. F is [ f/(1 f )] ln f, where f is the fraction of initial NO3 remaining at the time the culture is sampled. e is calculated to be 4.5% from the slope of the regression in (c), according to the Rayleigh integrated product equation (see text, eqn [5]). Dashed horizontal lines in (b) and (c) represent the initial d15N of the source NO3 (3.8%). Reprinted from Waser NAD, Harrison PJ, Nielsen B, Calvert SE, and Turpin DH (1998) Nitrogen isotope fractionation during the uptake and assimilation of nitrate, nitrite, ammonium, and urea by a marine diatom. Limnology and Oceanography 43: 215–224.
produced and nearly completely consumed within the open ocean surface mixed layer (Figure 3). Culture studies suggest an isotope effect for ammonium assimilation of up to B20%, decreasing as ammonium concentration decreases, and minimal isotope effects (o1%) for assimilation of nitrite and urea. In estuaries, where ammonium can accumulate in the shallow subsurface and be entrained into the surface layer, ammonium assimilation causes a clear increase in the d15N of the remaining ammonium pool. Isotope effect estimates based on ammonium concentrations and d15N in these environments range from 6.5% to 18.5% (Table 1).
Remineralization The return of organic N to nitrate occurs in two steps, the degradation of organic N to ammonium and the bacterial oxidation of ammonium to nitrate, or ‘nitrification’ (Figure 3). Nitrification itself occurs in two steps, the oxidation of ammonium to nitrite and the oxidation of nitrite to nitrate, mediated by distinct groups of microorganisms. Isotopic discrimination may occur at all steps involved in remineralization. Field studies generally suggest that both bacteria and zooplankton preferentially degrade low-d15N PN to ammonium, yielding residual organic matter
relatively high in 15N. The wide spectrum of reactions involved in organic N degradation and the heterogeneous nature of organic matter (comprised of compounds with distinct d15N that degrade at various rates) make quantifying the isotope effect associated with degradation difficult. A few laboratory studies have quantified the isotope effects of individual processes such as thermal peptide bond cleavage, bacterial amino acid uptake and transamination, and zooplankton ammonium release. Laboratory studies attempting to mimic degradation as a whole suggest a net isotope effect of r3%. Culture studies indicate a large isotope effect for the conversion of ammonium to nitrite, the first step in nitrification (Table 1). Estimates of the isotope effect for marine ammonia-oxidizing bacteria (B14–19%) are lower than those for terrestrial ammonia-oxidizing bacteria (as high as B38%), possibly due to phylogenetic differences. The isotope effect of nitrification estimated from ammonium concentration and d15N measurements in the Chesapeake Bay is 12–16%, similar to the culture results for marine ammonia-oxidizing bacteria (Table 1). Thus far, no culture-based information is available regarding isotope effects for ammonia-oxidizing crenarchaea or nitrite-oxidizing bacteria.
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46
NITROGEN ISOTOPES IN THE OCEAN
20
Nitrogen Reservoirs Dissolved Nitrogen
15 15NO3− (‰ vs. air)
Nitrate Nitrate accounts for most of the fixed N in the ocean. The d15N of deep ocean nitrate is typically B5%. Regionally, the d15N of nitrate varies between 2% and 20% due to the effects of N2 fixation, nitrate assimilation, and denitrification (Figure 5). Nitrate d15N significantly lower than deep-ocean nitrate has been observed in the upper thermocline of the low-latitude oligotrophic ocean (Figure 6). This 15N depletion is most likely due to the oxidation of newly fixed N, which, as described above, has a d15N of c. 1%. Values higher than 5% result from discrimination associated with nitrate assimilation by phytoplankton at the ocean surface (Figure 7) or denitrification in oxygendeficient zones of the ocean interior (Figure 8). Nitrate assimilation by phytoplankton leads to elevated d15N of nitrate in regions of the ocean where nitrate is incompletely consumed in surface waters, such as the high-latitude, nutrient-rich regions of the Southern Ocean and the subarctic Pacific, and the low-latitude upwelling regions of the California Current and the Equatorial Pacific. In the surface waters of these regions, there is a strong correlation between the degree of nitrate consumption by phytoplankton and the d15N of the nitrate remaining in the water (Figure 7). However, while nitrate assimilation elevates the d15N of nitrate in the surface ocean and causes modest 15N enrichment in some newly formed thermocline waters, it does not appear to affect greatly the d15N of nitrate in the deep ocean. Below 2.5–3.0- km depth in the ocean, nitrate d15N is relatively constant at B5%, despite large inter-basin differences in nitrate concentration. The minimal degree of isotopic variation in the nitrate of the deep ocean is due to the fact that, in most surface waters, the nitrate supply from below is almost completely consumed by phytoplankton, such that the organic N exported from the surface ocean converges on the d15N of the nitrate supply. Because the sinking flux d15N is close to that of the nitrate supplied from the ocean interior, remineralization of the sinking flux in the ocean interior does not alter greatly the d15N of deep nitrate. In this respect, the oceanic cycling of N isotopes differs markedly from that of the carbon isotopes. Because water-column denitrification occurs in the subsurface and because it consumes only a fraction of the nitrate plus nitrite available, its isotope effect is more completely expressed in the d15N of subsurface nitrate. In denitrifying regions of the water column, the d15N of nitrate in the subsurface is commonly elevated to 15% or higher (Figure 8). The
Water-column denitrification ∼ 25‰
10
Nitrate uptake ∼ 5‰
5
Newly fixed N added 15N ~ −1‰
Sedimentary denitrification ∼ 0‰ 0 0
0.5
1
1.5
2
[NO3−] (factor of initial value) Figure 5 The effect of different marine N cycle processes on nitrate d15N and concentration, assuming an initial nitrate d15N of 5%. The trajectories are for reasonable estimates of the isotope effects, and they depend on the initial nitrate d15N as well as the relative amplitude of the changes in nitrate concentration (30% for each process in this figure). A solid arrow denotes a process that adds or removes fixed N from the ocean, while a dashed line denotes a component of the internal cycling of oceanic fixed N. The effects of these two types of processes can be distinguished in many cases by their effect on the concentration ratio of nitrate to phosphate in seawater. The actual impact of the different processes on the N isotopes varies with environment. For instance, if phytoplankton completely consume the available nitrate in a given environment, the isotope effect of nitrate uptake plays no major role in the d15N of the various N pools and fluxes; the effect of nitrate generation by organic matter degradation and nitrification, not shown here, will depend on this dynamic. Similarly, the lack of a large isotope effect for sedimentary denitrification is due to the fact that nitrate consumption by this process can approach completion within sedimentary pore waters.
subsurface d15N maximum occurs in the core of the oxygen minimum and is correlated with the estimated degree of nitrate consumption by water-column denitrification. Denitrification, both in the water column and sediments, exerts a direct control on the d15N of mean deep ocean nitrate. When the ocean N budget is at steady state, the d15N of the fixed N removed (through water-column and sedimentary denitrification) will equal the d15N of the fixed N added (c. 1%, approximating N2 fixation as the sole source) (Figure 9). If denitrification with an isotope effect of 20–30% were occurring homogenously in the ocean water column and responsible for all fixed N loss, the d15N of mean oceanic nitrate would be 19–29% to achieve a d15N of –1% for N loss. That the modern mean oceanic nitrate d15N is B5%, much lower than 19–29%, reflects at least two factors: (1)
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NITROGEN ISOTOPES IN THE OCEAN
(a)
(b)
[N] (μM) 0
5
10
15
20
2
3
15N (‰ vs. air) 4
5
47
6
0 TON 15N
[TON]
Depth (m)
200
400
600
[NO3−]
NO3− 15N
800
1000
Figure 6 Depth profiles of (a) [NO3 ] (open circles) and [TON] (total organic N, or DON plus the small pool of PN) (open squares) and (b) nitrate d15N (filled circles) and TON d15N (filled squares) at the Bermuda Atlantic Time-Series Study site in the oligotrophic Sargasso Sea. Low nitrate d15N in the thermocline has been proposed to reflect N2 fixation at the surface, the N from which sinks and is remineralized at depth. The increase in nitrate d15N above 200 m reflects fractionation associated with nitrate assimilation. [TON] increases slightly into the surface layer while TON d15N decreases, suggesting a possible source of low-15N DON in the surface. The d15N of the TON pool is lower than that of mean ocean nitrate and deep nitrate at Bermuda (B5%) but higher than that of thermocline nitrate at Bermuda (B2.5%). High nitrate concentrations below 250 m prevent accurate assessment of TON d15N using published methods (see text). The d15N of sinking particles collected at 100 m (3.7%) is indicated by the arrow at top of (b); surface suspended PN d15N is c. 0.2% (not shown). Nitrate and TON data are the means of monthly measurements between June 2000 and May 2001. Modified from Knapp AH, Sigman DM, and Lipschultz F (2005) N isotopic composition of dissolved organic nitrogen and nitrate at the Bermuda Atlantic time-series study site. Global Biogeochemical Cycles 19: GB1018 (doi:10.1029/2004GB002320). Sinking and suspended PN d15N data from Altabet MA (1988) Variations in nitrogen isotopic composition between sinking and suspended particles: Implications for nitrogen cycling and particle transformation in the open ocean. Deep Sea Research 35: 535–554.
the importance of sedimentary denitrification, which appears to express a minimal isotope effect; and (2) the localized nature of water-column denitrification. With regard to the second, because denitrification consumes a significant fraction of the ambient nitrate in the ocean’s suboxic zones and elevates its d15N above that of the mean ocean (Figures 5 and 8), the d15N of nitrate being removed by water column denitrification is higher than if the substrate for denitrification had the mean ocean d15N. Much as with sedimentary denitrification, this reduces the expression of the organism-level isotope effect of water column denitrification and thus lowers the mean d15N of nitrate required to achieve an isotope balance between inputs and outputs. With these considerations, one study estimates that water-column denitrification is responsible for 30% of fixed N loss from the modern ocean, with sedimentary denitrification responsible for the remainder. Still, this isotope-based budget for marine fixed N remains uncertain. One limitation of using the N isotopes to investigate N cycling in the ocean is their inability to separate co-occurring processes with competing N isotopic signatures, such as denitrification/N2 fixation and nitrate assimilation/nitrification. Coupled
analysis of N and O isotopes in nitrate promises to disentangle such otherwise overprinting processes. Culture studies have demonstrated that the two most important nitrate-consuming processes, nitrate assimilation and denitrification, fractionate the N and O in nitrate with a ratio close to 1:1 (Figure 10). Deviations in the ratio of d18O and d15N in nitrate from 1:1 can therefore provide information about the nitrate being added by nitrification, such as if it derives from newly fixed N or if there is otherwise undetected nitrate cycling within a water parcel. Nitrite Nitrite is an intermediate in oxidative and reductive processes such as nitrification and denitrification. It also serves as a substrate for anammox and can be assimilated by phytoplankton. Only when these processes become uncoupled, such as at the base of the euphotic zone and in oxygen minimum zones, does nitrite accumulate to significant levels (0.1–10 mM). In some oxygen-deficient regions of the water column in the eastern tropical North Pacific, eastern tropical South Pacific, and the Arabian Sea, nitrite can represent as much as 25% of the pool of nitrogen oxides (NO3 þ NO2 ).
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48
NITROGEN ISOTOPES IN THE OCEAN
[NO3−] (μmol kg−1)
(a) 20
25
30
15N (‰ vs. air)
(b)
35
4
40
4.5
5
5.5
6
6.5
7
7.5
0 500 1000
Depth (m)
1500 2000 2500 3000 3500 4000 (c) 35
14 13
30
+
+
11 10
+ +
+ 15
+
+
9 ++ +
15NO3−
8
15N (‰ vs. air)
[NO3−] (μM)
25 20
12
[NO3−]
7
10
+ +
5
+
Sediment
+
+
15N
6 +
+
+ +
0
5 4
40
42
44
46
48
50
52
54
56
58
60
Latitude along 115° E (S) Figure 7 Nitrate d15N data from the Southern Ocean show the effect of uptake by phytoplankton. A depth profile of nitrate concentration (a) and d15N (b) from the Antarctic region of the east Indian Ocean (53.21 S, 1151 E) shows a decrease in nitrate concentration into the surface layer and an associated increase in nitrate d15N, both resulting from nitrate uptake by phytoplankton. A meridional transect (c) of nitrate concentration (open circles) and d15N (filled circles) in the surface mixed layer of the Southern Ocean along 115 1 E shows that nitrate d15N is spatially correlated with variations in the utilization of nitrate by phytoplankton, as reflected by the equatorward decrease in nitrate concentration. Nitrate d15N is lowest where the nitrate concentration is highest (and nitrate utilization is lowest), in Antarctic waters, south of B511 S in this region. The meridional gradient in nitrate d15N is recorded in bulk sediment d15N along 115 1 E (c, crosses) through the sinking of PN out of the surface ocean. Across the transect, sediment d15N is B5% higher than expectations for and measurements of sinking N d15N, probably mainly due to isotopic alteration of this N at the seafloor. Nevertheless, the link between nitrate consumption and sediment d15N provides a possible avenue for paleoceanographic reconstruction of nitrate utilization by phytoplankton. The left axis in (c) is scaled to 35 mM, the approximate nitrate concentration of Upper Circumpolar Deep Water that upwells in the Antarctic (see (a)). (a, b) Reprinted from Sigman DM, Altabet MA, Michener RH, McCorkle DC, Fry B, and Holmes RM (1997) Natural abundance-level measurement of the nitrogen isotopic composition of oceanic nitrate: An adaptation of the ammonia diffusion method. Marine Chemistry 57: 227–242. (c) Modified from Sigman DM, Altabet MA, McCorkle DC, Franc¸ois R, and Fischer G (2000) The d15N of nitrate in the Southern Ocean: Nitrate consumption in surface waters. Global Biogeochemical Cycles 13: 1149–1166, with the sediment data taken from Altabet MA and Franc¸ois R (1994) Sedimentary nitrogen isotopic ratio records surface ocean nitrate utilization. Global Biogeochemical Cycles 8: 103–116.
The d15N of nitrite is expected to reflect the balance of isotopic fractionation during nitrite production and consumption processes. In the eastern tropical North Pacific, the d15N of nitrite is very low
( 18% to 7%). This d15N is B30% lower than the nitrate in the same environment and is lower than expected from denitrification alone, give the known isotope effects for nitrate and nitrite reduction. The
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NITROGEN ISOTOPES IN THE OCEAN
15N of NO3−
(a)
6.0
0
9.0
Nitrate (μM)
(b) 12.0 15.0
500 Depth (m)
Depth (m)
0 10 20 30 40
0
500
1000
1500
1000
1500
2000
2000
0.2
49
0.4
0.6 15N
0
0.8
10 Nitrate deficit (μM)
of N2
Figure 8 (a) The d15N of NO3 (filled circles) and N2 (open circles) in water column profiles through an intense denitrification zone in the eastern tropical North Pacific (221 N, 1071 W). The shaded interval indicates the depth range with dissolved O2 concentration o10 mM where denitrification is encouraged (b, concentration shown in filled squares) and leads to a characteristic NO3 deficit (b, open squares) relative to phosphate. Measurements indicate enrichment of NO3 in 15N and concurrent depletion of N2 in 15N, arising from isotope discrimination during denitrification, with the conversion of NO3 to N2. e for denitrification in this environment was estimated to be B25%. Modified from Brandes JA, Devol AH, Yoshinari T, Jayakumar DA, and Naqvi SWA (1998) Isotopic composition of nitrate in the central Arabian Sea and eastern tropical North Pacific: A tracer for mixing and nitrogen cycles. Limnology and Oceanography 43: 1680–1689.
20 15 15N (‰ vs. air) 10 5
15N ~ 5‰
Sedimentary 30
0 −5
Denitrification
−10 −15
Water column
−20 Figure 9 Simplified global ocean N isotope budget. The y-axis indicates the d15N of a given flux or pool. The d15N of N from oceanic N2 fixation, the dominant N input to the ocean, is c. 1% (‘N2 fixation’ on the left). At steady state, the total denitrification loss (‘denitrification’ on the right) must have the same d15N as the input. The d15N of mean ocean nitrate is B5%. Water column denitrification removes nitrate with a low d15N (‘water column’ at lower right), while sedimentary denitrification removes nitrate with a d15N similar to that of mean ocean nitrate (‘sedimentary’ at upper right). The need for the flux-weighted d15N of the denitrification loss to be c. –1% leads to estimates of partitioning between water-column and sedimentary denitrification in which sedimentary denitrification is found to drive the greater part of the total N loss.
meaning of this low d15N is still not well understood but likely points to other processes acting on the NO2 pool. In these regions, nitrite can represent a significant reservoir of N that is depleted in 15N.
Δ18O of NO3− (‰ vs. initial)
N2 fixation
Mean ocean NO3−
18:15
= 1 0.9, 1.1
20
10
T. weissflogii T. pseudonana T. oceanica E. huxleyi
0 0
20 10 Δ15N of NO3− (‰ vs. initial)
30
Figure 10 The d18O vs. d15N in nitrate as it is progressively assimilated by four eukaryotic species of marine phytoplankton. Both d15N and d18O in nitrate increase as nitrate is consumed, and they do so with an O:N ratio for isotopic discrimination (18e:15e) of B1. Dashed lines show slopes of 1.1 and 0.9 for comparison. Modified from Granger J, Sigman DM, Needoba JA, and Harrison PJ (2004) Coupled nitrogen and oxygen isotope fractionation of nitrate during assimilation by cultures of marine phytoplankton. Limnology and Oceanography 49: 1763–1773.
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NITROGEN ISOTOPES IN THE OCEAN
Until recently, the two species nitrate and nitrite have typically been combined in isotopic analysis, such that the presence of low d15N nitrite may have masked some of the 15N enrichment of nitrate in the oxygen-deficient zone. Ammonium The d15N of ammonium reflects the production of ammonium by the degradation of organic N and its consumption by nitrification, ammonium assimilation, and perhaps anammox (Figure 3). Analytical constraints have limited isotopic studies of ammonium to environments with ammonium concentrations greater than 1 mM, excluding studies in the open ocean. In estuarine systems, where ammonium can be abundant, its d15N is often high (commonly higher than þ 10%, with one observation of þ 70%) and it increases as the ammonium concentration decreases along transects from riverine to marine waters, due to discrimination associated with ammonium consumption by nitrification and/or ammonium assimilation. In the open ocean interior, below the depth of algal assimilation, essentially all ammonium generated from particles is oxidized to nitrite and then nitrate before it can be transported into or out of a given region. Thus, nitrification should be of limited importance for the isotope dynamics of both particulate and dissolved N once the former has sunk out of the upper ocean. In the open ocean surface mixed layer, it is generally assumed that ammonium generated by remineralization is quickly and entirely assimilated by plankton, in which case the isotope effect associated with its consumption would not play an important role in N isotope dynamics of the open ocean. However, in at least some regions of the upper ocean, ammonium oxidation and assimilation are likely to co-occur. If the isotope effect of ammonium oxidation is greater than that of ammonium assimilation, low-d15N N will preferentially be routed to the nitrate pool by oxidation and high-d15N N will be routed back to the PN pool by assimilation. If the isotope effect of oxidation is less than that of assimilation, the opposite will occur. Thus, with better constaints on the isotope effects of ammoniumconsuming processes, the isotopes of upper ocean N pools promise to provide an integrative constraint on the relative rate of nitrification in the upper ocean, especially when paired with the O isotopes of nitrate (see above). Dissolved organic nitrogen DON concentrations are significant in the open ocean, typically Z4 mM in surface waters, decreasing to B2 mM in deep water. Fluxes associated with the DON pool are among
the least constrained terms in the modern marine N budget and may be important. Studies to date of bulk DON have been in the subtropical ocean, where DON is by far the dominant N pool in the surface ocean. In the surface mixed layer at the Bermuda Atlantic Time-series site in the Sargasso Sea, the concentration and d15N of TON are B4 mM and B4% (TON being total organic N, or DON plus the small pool of PN) (Figure 6). This d15N is similar to or slightly higher than the shallow subsurface nitrate that is entrained into the euphotic zone during wintertime vertical mixing (Figure 6(b)). Minimal gradients in the concentration and d15N of DON in this region of the upper ocean hinder reconstruction of fluxes of DON or the d15N of those fluxes. There is a weak increase in the concentration of TON into the surface layer and an accompanying decrease in its d15N (Figures 6(a) and 6(b)). Thus, there may be an input of low-d15N N into the surface DON pool, which is remineralized at depth, but this requires further validation. Progress on DON d15N dynamics would be aided by a method to remove nitrate from samples without compromising the DON pool, which would make subsurface waters and high-nitrate surface waters more accessible to study. Ongoing work on separable fractions of the DON pool (e.g., the highmolecular-weight fraction and its components) is also promising. Dissolved gases Dissolved N2 in equilibrium with atmospheric N2 at the surface has a d15N of 0.6%. The isotopic composition of dissolved N2 does not vary greatly in ocean profiles, except in zones of denitrification. Production of low-d15N N2 in denitrification zones results in measured N2 d15N as low as 0.2% (Figure 8(a)). Since N2 is the main product of denitrification, its d15N provides a test of the nitrate-based estimates of the isotope effect for this process. Dissolved N2O is produced by nitrification and is both produced and consumed by denitrification. The marine flux of N2O is perhaps one-third of the global flux of this greenhouse gas to the atmosphere; therefore, an understanding of the mechanisms of N2O production and their regulation in the ocean is an important goal. Culture studies indicate that bacterial production of N2O by nitrification and denitrification produces gas depleted in 15N and 18O relative to the source material. Consumption of N2O by denitrification leaves the residual gas enriched in 15 N and 18O, with d15N of N2O as high as 40% measured in the Arabian Sea. In oxygenated waters of the open ocean, nitrification likely dominates N2O production and its isotopic profile. A depth profile in
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NITROGEN ISOTOPES IN THE OCEAN
the subtropical North Pacific shows three main features (Figure 11): (1) isotopic equilibrium with atmospheric N2O at the surface; (2) a subsurface d15N minimum attributed to nitrification; and (3) a broad d15N maximum in deeper waters probably due to N2O consumption, perhaps in the denitrifying waters of the eastern Pacific margin. In and near denitrification zones, a strong maximum in the d15N of N2O is observed, presumably due to isotope fractionation associated with N2O consumption (via reduction to N2).
to surface waters. The isotopic effect of N recycling originates from heterotrophic processes. Zooplankton appear to release ammonium which has a lower d15N than their food source, making their tissues and solid wastes B3% higher in d15N than their food source. The low-d15N ammonium is consumed by phytoplankton and thus retained in the surface ocean N pool, while the 15N-enriched PN is preferentially exported as sinking particles, leading to a lower d15N of surface PN in regions where recycled N is an important component of the gross N supply to phytoplankton. Low d15N observed in suspended PN from the Antarctic and other highlatitude regions, beyond what is expected from isotope discrimination during nitrate assimilation, is unlikely to be due to N2 fixation and thus likely reflects N recycling. In the low-latitude, lownutrient ocean surface, such as the Sargasso Sea and western tropical Pacific, the relative importance of N2 fixation and N recycling in producing lowd15N surface particles is uncertain. Because of its implications for the rates of N2 fixation and N recycling, this question deserves further study. The d15N of suspended particles in the subsurface is typically B6% higher than suspended particles in the surface ocean and B3% higher than the sinking flux (Figure 12). The d15N of deep
Particulate Nitrogen
Suspended particles A typical profile of suspended particles has its lowest d15N in the surface layer, increasing below the euphotic zone (Figure 12). The d15N of suspended particles reflects in part the d15N of nitrate supplied to the surface ocean and, in nutrient-rich surface waters, isotope discrimination associated with its incomplete assimilation. However, the low d15N in the surface layer is typically lower than what would be expected solely from nitrate assimilation. This low d15N has two competing explanations: N2 fixation and N recycling. As described earlier, N2 fixation is expected to add fixed N with a d15N of c. 1% (a)
51
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Figure 11 Depth profiles of (a) N2O concentration and (b) d15N and (c) d18O of N2O at station ALOHA in the subtropical North Pacific (221 450 N, 1581 W) during four separate cruises. The solid line in (a) indicates theoretical saturation with atmospheric N2O at in situ temperatures and salinities. The minima in d15N and d18O around 200 m are thought to be due to significant in situ production of N2O from nitrification. The broad isotopic maxima at depth are likely due to N2O consumption, perhaps in the denitrifying waters along the eastern Pacific margin. The filled squares at the top of (b) and (c) represent measurements of d15N and d18O of atmospheric N2O during the Hawaii Ocean Time-series 76 cruise, and arrows indicate the range of historical measurements as of the late 1980s. Reprinted from Dore JE, Popp BN, Karl DM, and Sansone FJ (1998) A large source of atmospheric nitrous oxide from subtropical North Pacific surface waters. Nature 396: 63–66.
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15N (‰) −1
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1000 1500 2000 2500 3000 3500 Sediment trap Suspended particles Figure 12 Nitrogen isotopic values of suspended particulate matter and sinking particles (as collected by sediment traps) in the subtropical North Atlantic Ocean near BATS (311 500 N, 641 100 W; see Figure 6). The profiles of suspended PN show the representative depth gradient in d15N, with lower d15N in the surface ocean than at depth. The d15N of the sinking flux shows a decrease with depth, which has also been observed in other regions. Reprinted from Altabet MA, Deuser WG, Honjo S, and Stienen C (1991) Seasonal and depth-related changes in the source of sinking particles in the North Atlantic. Nature 354: 136–139.
particles is consistent with the inference that deep particles are the breakdown products of material exported from the surface, and that bacteria preferentially remineralize low-d15N PN. Isotopic analysis of zooplankton and organisms at higher trophic levels can provide insights into the marine N cycle. The ‘trophic effect’, an observed B3% increase in d15N per trophic level that presumably results from isotopic discrimination during metabolism of N-bearing organic matter, is used widely in food-web studies. N isotopic analysis of specific amino acids within organisms and organic matter promises new insights, as some amino acids increase in d15N with trophic level while others preserve the d15N of the food source. Sinking Particulate Nitrogen and sedimentary Nitrogen Because vertical sinking is an important mode of N export from the surface ocean, the d15N of the sinking flux is one of the most valuable N isotopic constraints on modern ocean processes. Combined with other isotopic data, sinking flux d15N data can provide information on the routes and mechanisms of nitrate supply and can be used to constrain other sources of N to the surface. The sinking flux also transfers the isotopic signal from
the surface ocean to the seafloor, providing the link through which the sediment column records the history of surface ocean processes (Figure 7(c)). Sinking particles collected in depth arrays of sediment traps often show a modest decrease in d15N with depth (Figure 12). This trend runs contrary to our expectations for the isotopic change of particulate matter as it degrades, and it currently lacks a compelling explanation. There is generally a good correlation between the d15N of surface sediments and sinking particulate d15N from the overlying water column. In regions of the ocean where a relatively large fraction of the organic rain is preserved in the sediment column, as occurs along continental margins, this correlation is excellent. In open ocean sediments where only a very small fraction of N is preserved, spatial patterns in the d15N of sediment core tops mirror those in the water column above (Figure 7(c)), but a significant 15 N enrichment (of B2–5%) is observed in the sediment N relative to sinking particles. Upon burial, reactions in the shallow sediment column known collectively as ‘diagenesis’ can cause a clear increase in the d15N of PN as it is incorporated into the sediment mixed layer. While some studies have found that sedimentary diagenesis has not greatly affected the paleoceanographic information provided by specific sedimentary records, it cannot be assumed that changes in the ‘diagenetic offset’ are unimportant. To address concerns regarding alteration of both sinking and sedimentary bulk d15N, studies are increasingly focusing on isolating specific N components, the d15N of which is insensitive to diagenesis, such as N bound within the mineral matrix of microfossils, or that does not change in d15N as it is degraded, such as chlorophyll degradation products. Nitrogen isotopes in the sedimentary record The isotopes of sedimentary N are used to investigate past changes in the marine N budget and the internal cycling of N within the ocean. The processes and parameters reflected by the d15N of sedimentary N include (1) mean ocean nitrate d15N, (2) regional subsurface nitrate 15N depletion or enrichment relative to the global ocean owing to N2 fixation or denitrification, (3) regional isotope dynamics associated with partial nitrate consumption in surface waters, and (4) possible direct contribution of newly fixed N to sinking PN. Paleoceanographers have focused on sediment d15N records underlying three environments where a single process or parameter is thought to dominate changes in sinking d15N. In oligotrophic regions, sediment d15N is assumed to reflect the d15N of mean ocean nitrate and therefore the global ocean
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NITROGEN ISOTOPES IN THE OCEAN
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Figure 13 Sedimentary d15N records spanning the past 140 ky, which encompass recent ice ages (shaded, marine oxygen isotope stages 2, 4, and 6, with 2 and 6 being the most extreme) and interglacials (stages 1, 3, and 5, with 1 and 5 being the most extreme). (a) Sediment record underlying the oligotrophic South China Sea (8 130.4 N, 112 119.9 E), where sedimentary d15N is expected to track, with some offset, the d15N of nitrate in the western Pacific thermocline. The western Pacific thermocline nitrate, in turn, is assumed to have maintained a constant isotopic relationship with deep ocean nitrate. The small magnitude of variation in d15N (o1.5%) and lack of correlation with glacial/interglacial transitions suggest that mean ocean nitrate d15N remained unchanged through shifts in Earth’s climate. (b) Sedimentary d15N record underlying the eastern tropical North Pacific (22 123.3 N, 107 104.5 W), a major region of denitrification. Interglacials are characterized by high d15N (8–9%), with d15N 2–3% lower during glacials. Low d15N, along with coincident evidence for decreased productivity and higher oxygen content in the mid-depth water-column, indicates decreased water column denitrification during glacial periods. (c) Sedimentary d15N record from the high-nitrate Antarctic Zone of the Southern Ocean (54 155 S, 73 150 E) shows higher d15N during the period spanning glacial stages 2–4, suggesting greater algal utilization of nitrate in the surface ocean. Coupled with evidence of lower glacial productivity, the glacial 15N enrichment suggests reduced nutrient supply from below. All data shown are of bulk sediment. The sedimentary d15N record shown in (b) is from a region of high organic matter preservation in the sediments, where bulk sedimentary d15N correlates well with sinking d15N. The records in (a) and (c) are from regions where a diagenetically driven difference is observed between sinking and sedimentary N, which introduces uncertainties in interpretation. (a) Core 17961 from Kienast M (2000) Unchanged nitrogen isotopic composition of organic matter in the South China Sea during the last climatic cycle: Global implications. Paleoceanography 15: 244–253. (b) Core NH8P from Ganeshram RS, Pederson TF, Calvert SE, and Murray JW (1995) Evidence from nitrogen isotopes for large changes in glacial–interglacial oceanic nutrient inventories. Nature 376: 755–758. (c) Core MD84-552 from Franc¸ois R, Altabet MA, Yu E-F, et al. (1997) Contribution of Southern Ocean surface-water stratification to low atmospheric CO2 concentrations during the last glacial period. Nature 389: 929–935.
balance of inputs and outputs of fixed N (Figure 13(a)). In denitrifying regions, sediment d15N has been taken to largely reflect changes in regional 15N enrichment due to water column denitrification (Figure 13(b)). In high-nutrient regions, sediment d15N primarily records the degree of nitrate consumption by algal assimilation, providing insight into changes in balance between gross nitrate supply to surface waters and export of organic N from the surface (Figure 13(c)). However, it must be kept in mind that multiple processes may affect the d15N of sediments in any given region. For example, even in low-nutrient regions far from denitrification zones, the sediment d15N record may record changes not only in mean ocean nitrate d15N
but also in the d15N of nitrate imported from subpolar regions and in the isotopic imprint of N2 fixation.
Concluding Remarks The study of the N isotopes in the ocean is young relative to that of the other light isotopes (e.g., carbon, oxygen, and sulfur), with much of the work to date developing the methods needed to measure different forms of oceanic N and establishing the isotope systematics of N cycle processes that are necessary to interpret observed patterns. Over the previous decades, the N isotopes have had perhaps their greatest impact on foodweb studies and in paleoceanographic
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NITROGEN ISOTOPES IN THE OCEAN
work. In the case of the latter, this reflects the ability of the N isotopes to provide basic constraints on environmental conditions when there are few other indicators available. Recent and ongoing method development is greatly improving our ability to measure diverse N pools in the ocean. This is yielding a new generation of N isotope studies that promise to provide geochemical estimates for the rates and distributions of N fluxes in the modern ocean, complementing instantaneous ‘bottle’ measurements of these fluxes as well as other geochemical approaches. Fundamental aspects of the oceanic N cycle are still poorly understood, and the N isotopes provide an important tool for their study.
See also Nitrogen Cycle. Redfield Ratio. Sedimentary Record, Reconstruction of Productivity from the.
Further Reading Altabet MA (2005) Isotopic tracers of the marine nitrogen cycle: Present and past. In: Hutzinger O (ed.) The Handbook of Environmental Chemistry, Vol. 2: Marine Organic Matter: Chemical and Biological Markers, pp. 251--294. Berlin: Springer. Altabet MA and Francois R (1994) The use of nitrogen isotopic ratio for reconstruction of past changes in surface ocean nutrient utilization. In: Zahn R, Kaminski M, Labeyrie L, and Pederson TF (eds.) Carbon Cycling in the Glacial Ocean: Constraints on the Ocean’s Role in Global Change, pp. 281--306. Berlin: Springer. Brandes JA and Devol AH (2002) A global marine fixed nitrogen isotopic budget: Implications for Holoene nitrogen cycling. Global Biogeochemical Cycles 16: 1120 (doi:10.1029/2001GB001856). Casciotti KL and McIlvin MR (2007) Isotopic analyses of nitrate and nitrite from references mixtures and application to Eastern Tropical North Pacific waters. Marine Chemistry 107: 184--201. Chang CCY, Silva SR, Kendall C, Michalski G, Casciotti KL, and Wankel S (2004) Preparation and Analysis of Nitrogen-bearing Compounds in Water for Stable Isotope Ratio Measurement. In: deGroot PA (ed.) Handbook of Stable Isotope Analytical Techniques, vol. 1, pp. 305--354. Amsterdam: Elsevier.
Fogel ML and Cifuentes LA (1993) Isotope fractionation during primary production. In: Engel MH and Macko SA (eds.) Organic Geochemistry, pp. 73--98. New York: Plenum. Fry B (2006) Stable Isotope Ecology. New York: Springer. Galbraith ED, Sigman DM, Robinson RS, and Pedersen TF (in press) Nitrogen in past marine environments. In: Bronk DA, Mulholland MR, and Capone DG (eds.) Nitrogen in the Marine Environment. Goericke R, Montoya JP, and Fry B (1994) Physiology of isotope fractionation in algae and cyanobacteria, In: Lajtha K and Michener R (eds.) Stable Isotopes in Ecology and Environmental Science, 1st edn., pp. 187--221. Oxford: Blackwell Scientific Publications. Macko SA, Engel MH, and Parker PL (1993) Early diagenesis of organic matter in sediments: Assessment of mechanisms and preservation by the use of isotopic molecular approaches. In: Engel MH and Macko SA (eds.) Organic Geochemistry, pp. 211--224. New York: Plenum. Michener R and Lajtha K (2007) Stable Isotopes in Ecology and Environmental Science, 2nd edn., New York: Wiley-Blackwell. Montoya JP (1994) Nitrogen isotope fractionation in the modern ocean: Implications for the sedimentary record. In: Zahn R, Kaminski M, Labeyrie L, and Pederson TF (eds.) Carbon Cycling in the Glacial Ocean: Constraints on the Ocean’s Role in Global Change, pp. 259--279. Berlin: Springer. Needoba JA, Sigman DM, and Harrison PJ (2004) The mechanism of isotope fractionation during algal nitrate assimilation as illuminated by the 15N/14N of intracellular nitrate. Journal of Phycology 40: 517--522. Owens NJP (1987) Natural variations in 15N in the marine environment. Advances in Marine Biology 24: 390--451. Peterson BJ and Fry B (1987) Stable isotopes in ecosystem studies. Annual Review of Ecology and Systematics 18: 293--320. Sigman DM, Granger J, DiFiore PJ, Lehmann MF, Ho R, Cane G, and van Geen A (2005) Coupled nitrogen and oxygen isotope measurements of nitrate along the eastern North Pacific margin. Global Biogeochemical Cycles 19: GB4022 (doi:10.1029/2005GB002458). Wada E (1980) Nitrogen isotope fractionation and its significance in biogeochemical processes occurring in marine environments. In: Goldberg ED, Horibe Y, and Saruhashi K (eds.) Isotope Marine Chemistry, pp. 375--398. Tokyo: Uchida Rokakudo.
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NOBLE GASES AND THE CRYOSPHERE M. Hood, Intergovernmental Oceanographic Commission, Paris, France Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1894–1898, & 2001, Elsevier Ltd.
Introduction Ice formation and melting strongly influence a wide range of water properties and processes, such as dissolved gas concentrations, exchange of gases between the atmosphere and the ocean, and dense water formation. As water freezes, salt and gases dissolved in the water are expelled from the growing ice lattice and become concentrated in the residual water. As a result of the increased salt content, this residual water becomes more dense than underlying waters and sinks to a level of neutral buoyancy, carrying with it the dissolved gas load. Dense water formation is one of the primary mechanisms by which atmospheric and surface water properties are transported into the interior and deep ocean, and observation of the effects of this process can answer fundamental questions about ocean circulation and the ocean–atmosphere cycling of biogeochemically important gases such as oxygen and carbon dioxide. Because it is not possible to determine exactly when and where dense water formation will occur, it is not an easy process to observe directly, and thus information about the rates of dense water formation and circulation is obtained largely through the observation of tracers. However, when dense water formation is triggered by ice formation, interaction of surface water properties with the ice and the lack of full equilibration between the atmosphere and the water beneath the growing ice can significantly modify the concentrations of the tracers in ways that are not yet fully understood. Consequently, the information provided by tracers in these ice formation areas is often ambiguous. A suite of three noble gases, helium, neon, and argon, have the potential to be excellent tracers in the marine cryosphere, providing new information about the interactions of dissolved gases and ice, the cycling of gases between the atmosphere and ocean, and mixing and circulation pathways in high latitude regions of the world’s oceans and marginal seas. The physical chemistry properties of these three gases span a wide range of values, and these differences cause them to respond to varying degrees to physical
processes such as ice formation and melting or the transfer of gas between the water and air. By observing the changes of the three tracers as they respond to these processes, it is possible to quantify the effect the process has on the gases as a function of the physical chemistry of the gases. Subsequently, this ‘template’ of behavior can be used to determine the physical response of any gas to the process, using known information about the physical chemistry of the gas. Although this tracer technique is still being developed, results from laboratory experiments and field programs have demonstrated the exciting potential of the nobel gases to provide unique, quantitative information on a range of processes that it is not possible to obtain using conventional tracers.
Noble Gases in the Marine Environment The noble gases are naturally occurring gases found in the atmosphere. Table 1 shows the abundance of the noble gases in the atmosphere as a percentage of the total air composition, and the concentrations of the gases in surface sea water when in equilibrium with the atmosphere. Other sources of these gases in sea water include the radioactive decay of uranium and thorium to helium-4 (4He), and the radioactive decay of potassium (40K) to argon (40Ar). For most areas of the surface ocean, these radiogenic sources of the noble gases are negligible, and thus the only significant source for these gases is the atmosphere. The noble gases are biogeochemically inert and are not altered through chemical or biological reactions, making them considerably easier to trace and quantify as they move through a system than other gases whose concentrations are modified through reactions. The behavior of the noble gases is largely determined by the size of the molecule of each gas and the natural affinity of each gas to reside in a Table 1
Noble gases in the atmosphere and sea water
Gas
Abundance in the atmosphere (%)
Concentration in seawater (cm3 g 1)
Helium Neon Argon
0.0005 0.002 0.9
3.75 108 1.53 107 2.49 104
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gaseous or liquid state. The main physical chemistry parameters of interest are the solubility of the gas in liquid, the temperature dependence of this solubility, and the molecular diffusivity of the gas. The suite of noble gases have a broad range of these properties, and the behavior of the noble gases determined by these properties, can serve as a model for the behavior of most other gases. One unique characteristic of the noble gases that makes them ideally suited as tracers of the interactions between gases and ice is that helium and neon are soluble in ice as well as in liquids. It has been recognized since the mid-1960s that helium and neon, and possibly hydrogen, should be soluble in ice because of the small size of the molecules, whereas gases having larger atomic radii are unable to reside in the ice lattice. These findings, however, were based on theoretical treatises and carefully controlled laboratory studies in idealized conditions. It was not until the mid-1980s that this process was shown to occur on observable scales in nature, when anomalies in the concentrations of helium and neon were observed in the Arctic. The solubility of gases in ice can be described by the same principles governing solubility of gases in liquids. Solubility of gases in liquids or ice occurs to establish equilibrium, where the affinities of the gas to reside in the gaseous, liquid, and solid state are balanced. The solubility process can be described by two principle mechanisms: 1. creation of a cavity in the solvent large enough to accommodate a solute molecule; 2. introduction of the solute molecule into the liquid or solid surface through the cavity. In applying this approach to the solubility of gases in ice, it follows that if the atomic radius of the solute gas molecule is smaller than the cavities naturally present in the lattice structure of ice, then the energy required to make a cavity in the solvent is zero, and the energy required for the solubility process is then only a function of the energy required to introduce the solute molecule into the cavity. For this reason, the solubility of a gas molecule capable of fitting in the ice lattice is greater than its solubility in a liquid. The solubilities of helium and neon in ice have been determined in two separate laboratory studies, and although the values agree for the solubility of helium in ice, the values for neon disagree. The size of neon is very similar to the size of a cavity in the ice lattice, and the discrepancies between the two reported values for the solubility of neon in ice may result from small differences in the experimental procedure. During ice formation, most gases partition between the water and air phases to try to establish
Table 2
Noble gas partitioning in three phases
Partition phases
Helium
Neon
Argon
Bubble to water Bubble to ice Ice to water
106.8 56.9 1.9
81.0 90.0, 56.3 0.9, 1.4
18.7 N 0
equilibrium under the changing conditions, whereas helium and neon additionally partition into the ice phase. As water freezes, salt and gases are rejected from the growing ice lattice, increasing the concentrations of salt and gas in the residual water. Helium and neon partition between the water and ice reservoirs according to their solubility in water and ice. The concentrations of the gases in the residual water that have been expelled from the ice lattice, predominantly oxygen and nitrogen, can become so elevated through this process that the pressure of the dissolved gases in the water exceeds the in situ hydrostatic pressure and gas bubbles form. The gases then partition between the water, the gas bubble, and the ice according to the solubilities of the gases in each phase. This three-phase partitioning process can occur either at the edge of the growing ice sheet at the ice–water interface, or in small liquid water pockets, called ‘brine pockets’ in salt water systems, entrained in the ice during rapid ice formation. Table 2 quantitatively describes how the noble gases partition between the three phases when a system containing these three phases is in equilibrium in fresh water at 01C. The numbers represent the amount of the gas found in one phase relative to the other. For example, the first row describes the amount of each gas that would reside in the gaseous bubble phase relative to the liquid phase; thus for helium, there would be 106.8 times more helium present in the bubble than in the water. This illustrates the small solubility of helium in water and its strong affinity for the gas phase. Because helium is 1.9 times more soluble in ice than in water, helium partitions less strongly between the bubble and ice phases compared to the partition between the bubble and water phases. The two numbers shown for neon represent the two different estimates for the solubility of neon in the ice phase. One estimate suggests that neon is less soluble in ice than in water, whereas the other suggests that it is more soluble in ice.
Application of the Noble Gases as Tracers The noble gases have been used as tracers of air–sea gas exchange processes for more than 20 years.
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NOBLE GASES AND THE CRYOSPHERE
Typically, the noble gases are observed over time at a single location in the ocean along with other meteorological and hydrodynamic parameters to characterize and quantify the behavior of each of the gases in response to the driving forces of gas exchange such as water temperature, wind speed, wave characteristics, and bubbles injected from breaking waves. Because both the amount and rate of a gas transferred between the atmosphere and ocean depend on the solubility and diffusivities of the gas, the noble gases have long been recognized as ideal tracers for these processes. In addition, argon and oxygen have very similar molecular diffusivities and solubilities, making argon an excellent tracer of the physical behavior of oxygen. By comparing the relative concentration changes of argon and oxygen over time, it is possible to account for the relative contributions of physical and biological processes (such as photosynthesis by phytoplankton in the surface ocean) to the overall concentrations, thus constraining the biological signal and allowing for estimates of the biological productivity of the surface ocean. The observations of anomalous helium and neon concentrations in ice formation areas and the suggestion that these anomalies could be the result of solubility of these gases in the ice were made in 1983, and since that time, a number of laboratory and field studies have been conducted to characterize and quantify these interactions. The partitioning of the noble gases among the three phases of gas, water, and ice creates a very distinctive ‘signature’ of the noble gas concentrations left behind in the residual water. Noble gas concentrations are typically expressed in terms of ‘saturation’, which is the concentration of a gas dissolved in the water relative to its equilibrium with the atmosphere at a given temperature. For example, a parcel of water at standard temperature and pressure containing the concentrations of noble gases shown in column 2 of Table 1 would be said to have a saturation of 100%. Saturations that deviate from this 100% can arise when equilibration with the atmosphere is incomplete, either because the equilibration process is slow relative to some other dynamic process acting on the system (for example, rapid heating or cooling, or injection of bubbles from breaking waves), or because full equilibration between the water and atmosphere is prevented, as in the case of ice formation. Typical saturations for the noble gases in the surface ocean range from 100 to 110% of atmospheric equilibrium, due mostly to the influx of gas from bubbles. Ice formation, however, can lead to quite striking saturations of 70 to 60% for helium and neon and þ 230% for argon in the relatively
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undiluted residual water. Ice melting can also lead to large anomalous saturations of the noble gases, showing the reverse of the freezing pattern for the gas saturations, where helium and neon are supersaturated while argon is undersaturated with respect to the atmosphere. The interactions of noble gases and ice have been well-documented and quantified in relatively simple freshwater systems. Observations of large noble gas anomalies in a permanently ice-covered antarctic lake were quantitatively explained using the current understanding of the solubility of helium and neon in ice and the partitioning of the gases in a three-phase system. Characteristics of ice formed from salt water are more complex than ice formed from fresh water, and the modeling of the system more complex. Using a set of equations developed in 1983 and measurements of the ice temperature, salinity, and density, it is possible to calculate the volume of the brine pockets in the ice and the volume of bubbles in the ice. With this type of information, a model of the ice and the dissolved gas balance in the various phases in the ice and residual water can be constructed. Such an ice model was developed during a field study of gas–ice interactions in a seasonally ice-covered lagoon, and the model predicted the amount of argon, nitrogen, and oxygen measured in bubbles in similar types of sea ice. No measurements are available for the amount of helium and neon in the bubbles of sea ice to verify the results for these gases. It is also possible to predict the relative saturations of the noble gases in the undiluted residual water at the ice– water interface, and this unique fingerprint of the noble gases can then serve as a tracer of the mixing and circulation of this water parcel as it leaves the surface and enters the interior and deep ocean. In this manner, the supersaturations of helium from meltwater have been successfully used as a tracer of water mass mixing and circulation in the Antarctic, and the estimated sensitivity of helium as a tracer for these processes is similar to the use of the conventional tracer, salinity, for these processes. As an illustration of the ways in which the noble gases can be used to distinguish between the effects of ice formation, melting, injection of air bubbles from breaking waves, or temperature changes on dissolved gases, Figure 1 shows a vector diagram of the characteristic changes of helium compared to argon resulting from each of these processes. From a starting point of equilibrium with the atmosphere (100% saturation), both helium and argon saturations increase as a result of bubbles injected from breaking waves. Ice formation increases the saturation of argon and decreases the saturation of helium, whereas ice melting has the opposite effect.
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NOBLE GASES AND THE CRYOSPHERE 114
Helium saturation (%)
112 110
Air injection
Melting
108 106 104
mass was formed and the transformations it has experience since leaving the surface ocean. These issues are important for our understanding of the global cycling of gases between the atmosphere and the ocean and for revealing the circulation pathways of water in the Arctic, Antartic, and high latitude marginal seas. The noble gases could represent a significant addition to the set of tracers typically used to study these processes.
102 100
Rapid cooling
98 97 97.5
Freezing
See also 98 98.5 99
99.5 100 100.5 101 101.5 102
Argon saturation (%)
Figure 1 Vector diagram of helium and argon saturation changes in response to upper ocean processes.
Changes in temperature with no gas exchange with the atmosphere to balance this change can lead to modest changes in the saturations of the gases, where the gas saturations decrease with decreasing temperature and increase with increasing temperature. The trends presented here are largely qualitative indicators, since quantitative assessment of the changes depend on the exact nature of the system being studied. However, this diagram does illustrate the general magnitude of the changes that these processes have on the noble gases and conversely, the ability of the noble gases to differentiate between these effects.
Conclusions The use of the noble gases as tracers in the marine cryosphere is in its infancy. Our understanding of the interactions of the noble gases and ice have progressed from controlled, idealized laboratory conditions to natural freshwater systems and simple salt water systems, and the initial results from these studies are extremely encouraging. This technique is currently being developed more fully to provide quantitative information about the interactions of dissolved gases and ice, and to utilize the resulting effects of these interactions to trace water mass mixing and circulation in the range of dynamic ice formation environments. Water masses in the interior and deep ocean originating in ice formation and melting areas have been shown to have distinct noble gas ratios, which are largely imparted to the water mass at the time of its formation in the surface ocean. By understanding and quantifying the processes responsible for these distinct ratios, we will be able to learn much about where and how the water
Air–Sea Gas Exchange. Arctic Ocean Circulation. Bottom Water Formation. Bubbles. CFCs in the Ocean. Ice–ocean interaction. Long-Term Tracer Changes. Oxygen Isotopes in the Ocean. Polynyas. Sea Ice: Overview. Stable Carbon Isotope Variations in the Ocean. Sub Ice-Shelf Circulation and Processes. Tritium–Helium Dating. Water Types and Water Masses.
Further Reading Bieri RH (1971) Dissolved noble gases in marine waters. Earth and Planetary Science Letters 10: 329--333. Cox GFN and Weeks WF (1982) Equations for determining the gas and brine volumes in sea ice samples, USA Cold Regions Research and Engineering Laboratory Report 82-30, Hanover, New Hampshire. Craig H and Hayward T (1987) Oxygen supersaturations in the ocean: biological vs. physical contributions. Science 235: 199--202. Hood EM, Howes BL, and Jenkins WJ (1998) Dissolved gas dynamics in perennially ice-covered Lake Fryxell, Antarctica. Limnology and Oceanography 43(2): 265--272. Hood EM (1998) Characterization of Air–sea Gas Exchange Processes and Dissolved Gas/ice Interactions Using Noble Gases. PhD thesis, MIT/WHOI, 98–101. Kahane A, Klinger J, and Philippe M (1969) Dopage selectif de la glace monocristalline avec de l’helium et du neon. Solid State Communications 7: 1055--1056. Namoit A and Bukhgalter EB (1965) Clathrates formed by gases in ice. Journal of Structural Chemistry 6: 911--912. Schlosser P (1986) Helium: a new tracer in Antarctic oceanography. Nature 321: 233--235. Schlosser P, Bayer R, Flodvik A, et al. (1990) Oxygen-18 and helium as tracers of ice shelf water and water/ice interaction in the Weddell Sea. Journal of Geophysical Research 95: 3253--3263. Top Z, Martin S, and Becker P (1988) A laboratory study of dissolved noble gas anomaly due to ice formation. Geophysical Research Letters 15: 796--799. Top Z, Clarke WB, and Moore RM (1983) Anomalous neon–helium ratios in the Arctic Ocean. Geophysical Research Letters 10: 1168--1171.
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NON-ROTATING GRAVITY CURRENTS P. G. Baines, CSIRO Atmospheric Research, Aspendale, VIC, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1898–1903, & 2001, Elsevier Ltd.
Gravity currents (also known as ‘density currents’) in the ocean are flows of water that are principally due to differences in density (due to differing temperature and/or salinity) between neighboring water bodies. These density differences cause lateral pressure gradients that produce the horizontal motion. They may occur on length scales ranging from centimeters to hundreds of kilometers, and they have a characteristic structure that is described below. In the oceanic environment there are three main types, as follows. Firstly, gravity currents may move along the ocean bottom, with denser (colder or saltier) water moving under a lighter water mass. Secondly, a body of lighter water may move along the ocean surface, above denser water, and thirdly, a homogeneous body of water of intermediate density may penetrate a larger stratified water body of varying density, with lighter fluid above and denser fluid below. The latter process is termed an intrusion. If these flows last for more than a significant fraction of the inertial period, the Earth’s rotation (via the Coriolis force) has an important effect on the flow, and these rotational effects are discussed elsewhere (see Rotating Gravity Currents). This article gives some examples of these flows, and then describes their basic dynamical properties for flows that move horizontally in one direction, and that move radially outward in two dimensions. It then proceeds to discuss gravity currents that flow down slopes, taking into account the effects of environmental stratification.
Examples Gravity currents exist in the ocean in a wide variety of forms. On the smallest scale of interest here, they are man-made. Prominent examples are sewage outfalls, in which effluent is piped to some distance offshore and is then released into the ocean. If the effluent is denser than the environment it may spread laterally over the local bottom, but if lighter, it may rise in a plume to the surface and spread there. Another example is effluent from power stations; this may be released on the surface, but it then behaves in a similar fashion to the sewage outfall. Oil released
on the surface (as from a shipwrecked tanker) will initially spread as a buoyant gravity current, although being immiscible with water, its mixing properties will be quite different from the previous examples. Moving to larger scale, gravity currents may be found in rivers, lakes, and estuaries. A sudden surge of water down a river into a lake or estuary can form a gravity current on the bottom or at the surface. If the lake is deep and density-stratified, and the flow into the lake is denser (e.g., colder) than the surface water, most of this inflow may begin as a downslope current, but end up as an intrusion. Several examples in the coastal environment are due to tides, which can cause the shoreward flow of buoyant fluid into a region on a broad front, or alternatively, a salt wedge of dense fluid moving inward on the bottom. In the open ocean, fronts (see Shelf Sea and Shelf Slope Fronts) that separate water of different densities can exist for periods of days, weeks, and months in much the same location. These fronts are maintained in dynamical balance by the Coriolis force, but they may exhibit gravity current character locally at times when the flow is disturbed and the geostrophic balance is broken. Sometimes turbulence within a current will cause the sediment on the bottom to become suspended in the water. When this occurs, the sediment can add to the fluid density, and the result constitutes a turbidity current. This may cause a redistribution of sediment, often to deeper water. Earthquakes may trigger submarine landslides that transport and redistribute sediment over large distances. Over timescales of millions of years, these may become conspicuous layers of sedimentary rock (turbidites) that leave a notable signature in the geological record.
Unidirectional Density Currents The dynamical essentials of the above phenomena can be encapsulated by looking at simple idealized flows that represent them. These can best be seen from simple laboratory experiments. A density current can be created in the laboratory by releasing cold water into relatively warmer water, but it is usually simpler to work instead with isothermal fresh water, and use dissolved salt to make the water denser. In Figure 1 a body of dense, dyed salty water has been released into a tank of fresh water, producing a density current moving from left to right over a horizontal surface, viewed from the side. The leading part of the current consists of a ‘head’ with an overhanging leading nose. This head may be
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In order to quantify this process, an idealized twodimensional model of a density current can be considered, in which the volume flux Q and density r1 of the dense fluid are specified at a given source location. The Reynolds number (¼ Q=n where n is kinematic viscosity) is assumed to be large. If the density of the environmental fluid is r0 , the buoyancy of the dense fluid is then g0 ¼ ðr1 r0 Þg=r0 Behind the head, the velocity v1 and thickness d1 of the dense layer are approximately constant. We then have Q ¼ v1 d1 , and from dimensional analysis alone we have: d1 BðQ2 =g0 Þ1=3 Figure 1 A vertical cross-section showing the structure of the flow in and behind the head of a gravity current advancing over a rigid bottom in a laboratory experiment, from left to right. The dense fluid has been marked by fluorescein dye. Note the dense fluid layer on bottom at the left, and the mixed layer above it. (Dyed salt water inflow density 1.05 g cm3 into fresh water, current depth of order 1–2 cm.)
regarded as a limiting form of hydraulic jump in a cold layer of dense fluid (see Overflows and Cascades), in which the depth upstream of the jump is zero. Behind the head the fluid has a three-layer structure, of which the first layer of dense fluid constitutes the main part of the current, at the bottom. This fluid moves faster than the head, and catches up with it. Inside the head it rises and is mixed with the surrounding lighter environmental fluid, and forms a density-stratified layer, spread out above the bottom current. Above this is the third layer of unmixed environmental fluid. Vorticity is produced in the upper part of the head by shear between the head and the ambient fluid, and this is also deposited in the stratified layer. This mixed stratified layer moves slowly in the direction of the current. It is highly turbulent immediately behind the head, where most of the mixing takes place, and this turbulence decays with distance from the head. Here the interfaces between the mixed layer and the dense fluid below and the environmental fluid above are generally stable, so that apart from the decaying turbulence little further significant mixing occurs. Density current heads have three-dimensional structure containing many lumps and bumps, and these shapes change continually as the head propagates. Some small parts of the head are more advanced than others, but these then disappear and are overtaken by others. This lobe and cleft structure is due to the drag on the current by the rigid lower surface over which it propagates. This retards the lowest levels of dense fluid, causing the overhanging noses. The lighter fluid beneath these noses then causes the unstable three-dimensional lobe and cleft structure.
v1 BðQg0 Þ1=3 Bðg0 d1 Þ1=2
½1
The speed vH of the head also scales with v1, with vH ov1 . Laboratory observations of flow into a deep homogeneous environment at large Reynolds numbers give: vH ¼ 1:2ðg0 d1 Þ1=2 v1 ¼ 1:4ðg0 d1 Þ1=2
½2
which are approximately constant with time for a sustained source of fluid. The same observations also show that the total height of the head is typically 3d1, and the height of the nose is about 0.15d1. These expressions apply to density current heads in general, when the buoyancy and thickness of the dense fluid approaching the head are given by g0 and d1. This steady pattern is maintained because the driving pressure gradient is retarded by loss of momentum through mixing in the head, and the following current is restrained by friction with the bottom surface and the fluid above. When the current becomes long, the thickness of the dense layer decreases gradually away from the source, providing a small pressure gradient to overcome friction. If buoyant fluid is released into denser fluid causing a gravity current that flows along the surface, the lack of stress on the surface implies that the leading nose is smooth rather than overhanging, and the lobe and cleft structure is largely absent (see Figure 2). Experiments show that these flows may be significantly affected by surface tension unless the latter is very small compared with buoyancy, but where buoyancy dominates (as in most cases in the ocean) the above dynamical relationships in eqn [2] apply with slightly different constants. Unidirectional Density Currents due to Collapse of a Dense Body of Fluid
Density currents may also be created by suddenly releasing a large body of dense fluid within a stationary environment of lighter fluid. This models a finite-sized source, and is readily simulated in a laboratory tank by raising a vertical barrier that
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NON-ROTATING GRAVITY CURRENTS
61
< L0 > <
>
D0
x
Figure 2 As for Figure 1, but showing a section along a gravity current advancing to the left beneath a free surface. Note the (relative) absence of an overhanging nose. (Dyed freshwater inflow salt water of density 1.1 g cm3, current depth of order 1– 2 cm.)
separates the dense and lighter fluid. In a twodimensional situation where the fluid spreads in the x-direction only, the flow passes successively through three stages: the slumping, inertia–buoyancy, and viscous stages. In the first stage of collapse of the dense fluid (of initial depth d0, length L0, and area A ¼ A0 ¼ d0 L0 in fluid of total depth D0), termed the slumping stage, the front travels as a density current of constant speed and depth. However, the surface of the dense fluid behind this front is not horizontal, as large amplitude internal waves propagate on it (see Figure 3). These waves reflect from the left-hand end (or center, for a symmetric collapsing body), and the character of the flow changes when they catch up to the front, or density current head. If the total depth is sufficiently small, or if the upper level fluid is sufficiently strongly stratified to (partially) simulate a rigid lid, the resulting motion of the ambient fluid may affect the waves on the interface with the dense fluid, and hence affect the details of the slumping behavior. This effect may be represented in experiments by a finite total depth, as shown schematically in Figure 3. In particular, the presence of a reversed flow in the upper layer causes a corresponding reduction in the speed over the ground of the collapsing front, relative to eqn [2]. The distance xs at which the slumping stage ends and the waves reach the head, measured from the end of the tank, is then observed to be: xs =L0 ¼ 3 þ 7:4d0 =D0
½3
After the reflected waves have caught up with the density current head the flow enters the second selfsimilar stage, which is dominated by inertia,
Figure 3 Schematic diagrams showing various stages at successive times in the initial slumping phase of the collapse of a rectangular volume (shown dashed) of dense fluid with depth d0 equal to the total depth D0. The mixed region has been omitted from this sketch.
buoyancy, and mixing. Here the dense fluid collapses in the form of a rectangle of approximately uniform area, with increasing length and uniformly decreasing height d1. This rectangular uniformity is maintained by internal waves propagating outwards on the dense fluid interface. Mixing in the main body of the collapsing rectangular current is generally very small. This is because the mean gradient Richardson number at its upper boundary (on which mixing depends) is generally 40.25, implying that the mean flow is stable and does not generate local mixing (see Upper Ocean Mixing Processes). Hence the density within the collapsing rectangular body of dense fluid remains largely unaltered until the fluid enters the head, where most of the mixing takes place. The total rectangular area AðtÞ of unmixed dense fluid continually decreases because v1 > vH in eqn [2]. During this stage, the length of the dense fluid increases as ðAg0 t2 Þ1=3 , so that the velocity of the density current head decreases with time as ðAg0 t2 Þ1=3 because of the steady decrease of d1. This continues until the third stage is reached, where the dense layer becomes sufficiently thin and slow for viscous effects to become important, a regime that is not relevant here. In the second stage, dense fluid enters the head from behind where it is mixed with environmental fluid. The resulting mixture is spread out behind, in a layer over the dense layer. The rate of mixing within the head is roughly proportional to its mean height, and hence decreases with it. At early times in this stage ðx > xs Þ, the detrained mixed fluid consists mostly of the dense fluid with a small part of environmental fluid. However, as the current proceeds, this proportion reverses and near the end (ðxcxs ), nearly all the mixed fluid is environmental. When the end of the second stage has been reached (at 1=2 x ¼ xs þ 29A0 ), and mixing has effectively ceased, the total volume of mixed fluid that has been
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produced is slightly more than twice the initial volume of dense fluid. The volume of the remaining unmixed dense fluid is quite small, so that overall, the dense fluid and environmental fluid have been mixed in approximately equal proportions.
Radial Density Currents Collapsing localized bodies of dense fluid may be constrained to spread in one direction if they are spatially confined, such as in a submarine valley, but often they are free to spread horizontally without confinement. These have a curved front or head, expanding radially away from a central source. Provided that the curvature of the front is not too large, the speed of the front and of the fluid behind it are given approximately by eqn [2]. Such flows may be modeled by axisymmetric collapsing bodies of dense fluid, which have some properties that resemble those of one-dimensional spreading. Again, there are three stages of spreading: the slumping, inertia–buoyancy balance, and viscous stages. For a body of dense fluid of initial height d0, radius R0, and volume V0, in fluid of overall depth D0, the initial collapse occurs in the ‘slumping’ stage, governed by wave propagation and adverse flow of the ambient fluid. Here the radial position R of the front increases at a constant speed in the range R0oRoRs, where Rs corresponds to xs and denotes the limit of the slumping stage. The shear between the front of dense fluid and the ambient fluid produces vorticity that is initially contained within the mixing head of the current, and as the front expands this vorticity increases by vortex stretching. This initial vortex intensifies as it expands, and may be much deeper than the following fluid. For a small initial volume of dense fluid, this expanding vortex ring of dense mixed fluid is all that is produced. But, for a larger initial volume, as the dense fluid expands and the head moves forward, the initial vortex progressively breaks up and is subsumed into the mixed layer. New vorticity is continually created at the head, as for the unidirectional currents, but this is associated with newly mixed fluid and the stretching is much less than that in the initial vortex. In this second inertial–buoyancy phase (where R>Rs), gravity waves maintain approximately uniform depth of the body of dense fluid, and it collapses in the form of an axisymmetric pillbox, with the volume slowly decreasing due to mixing and detrainment behind the circular head. If the volume of dense fluid at any given time is V ¼ pR2 d1 , and the properties of the front are given by eqn [2], the radius R increases as: RBðg0 VÞ1=4 t1=2
½4
V slowly decreases as dense fluid entering the head is mixed, and this dynamical regime continues until (in the laboratory) the viscous regime is reached. If the source of dense fluid is maintained with constant volume flux Q1, the initial vortex and head form as above, but the flow behind it evolves differently, as the depth d1 and velocity v1 of the following radial outflow are not uniform and decrease with radial distance r. The only significant mixing occurs behind the head, and for steady flow, dimensional analysis gives: 2 1=3 Q Q1 g0 1=3 v1 B ½5 d1 B 0 12 gr r In the inertio-gravity range, the radial speed of the head is then given by: Q1 g0 1=4 ½6 vH ¼ 0:63 t
Density Currents Down Slopes When the terrain is not horizontal but slopes downwards at angles greater than about 51, buoyancy acts directly to drive the current downward, and this provides stronger forcing than the indirect effect of establishing a horizontal pressure gradient as described above. The downslope buoyancy force is now g0 siny per unit mass of dense fluid, where y is the angle between the slope of the terrain and the horizontal. The onset of a steady source of dense fluid at the top of the slope leads to the formation of a density current head similar to that described above. In the current following the head the physics is different because the flow is now unstable, unless the slope angle is very small, and mixing now occurs along the whole length of the density current, not just behind the head. After the passage of the head an approximately steady flow is established, with the buoyancy force being balanced primarily by the entrainment of environmental fluid into the current from above and, to a much lesser extent, by the drag on the bottom surface. For a homogeneous environment, both the size of the head and the thickness of the following current increase with downslope distance. Behavior of the Head
For a constant supply of dense fluid of buoyancy g00 and flow rate Q0 at the top of the slope, the speed of the head vH is almost independent of slope angle and is given approximately by: vH ¼ ð1:570:2Þðg00 Q0 Þ1=3
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½7
NON-ROTATING GRAVITY CURRENTS
However, in spite of this uniform speed, the size of the head increases as it moves downslope, at the rate: dH=dx ¼ 0:72y=p
½8
where H is the height (or thickness) of the head, x is downslope distance, and y is in radians. The increased buoyancy at steeper slopes is balanced by increased entrainment that acts to keep the head speed constant, regardless of downslope distance and slope angle. Entrainment in Downslope Flows
Behind the head the flow is unstable, and mixing with the overlying fluid results. This process becomes stronger with increasing slope angle y. The mixed fluid above the dense layer is also denser than the environment, and hence it also moves downslope under gravity but at a reduced speed. The thickness and volume of this mixed layer both increase with downslope distance. This combined flow may now be regarded as a single entity, and the net downslope buoyancy flux is constant with x and t, and equal to g0 0Q0. The net entrainment of environmental fluid into this overall downslope flow may be described by an entrainment coefficient E which is a function of the bulk Richardson number Ri, defined by: Ri ¼ g0 d cosy=U2
½9
where U is the mean velocity and d the mean thickness of the total flow (see Upper Ocean Mixing Processes). At each level, the velocity of inflow we of the environmental fluid into the downflow is given by: we ¼ EðRi ÞU
½10
63
buoyancy frequency (see Upper Ocean Vertical Structure) of the ambient stratification, a depth D may be identified below the source, defined by N2 ¼ g0 0/D, where the ambient density equals the inflowing density. If there were no mixing, all of the inflowing fluid would be expected to reach and spread horizontally at this level. The speed of the head of a downslope gravity current into a stratified environment is again approximately constant over most of its distance travelled, and scales with eqn [7] above. The main differences from the homogeneous case concern the following current, which is discussed next. For slope angles less than about 201, which covers most cases of interest in the ocean, it is appropriate to regard the main dense current and the mixed layer above it as separate entities. In laboratory experiments a clear interface is visible between them, as seen in Figure 4, although there are turbulent fluxes across it. These flows are governed by two dimensionless parameters – a bulk Richardson number Ri defined as in eqn [9] but based on the dense layer only, and a parameter M, defined by: M ¼ QN3 =g02
½11
which is a measure of the effect of the ambient stratification. The dense layer is observed to have approximately uniform thickness over most of its length, but its velocity mostly decreases with downslope distance. It loses fluid to the mixed layer by detrainment, but also entrains fluid from it so that its density progressively decreases, and the fluid that remains in the dense current reaches its ambient level where it spreads out, at a height somewhere above the level of D below the source. Entrainment into this dense layer may be expressed in terms of an
where U is the mean velocity of the downflow. E decreases monotonically with increasing Ri, from E ¼ 0.075 at Ri ¼ 0 to very small values for Ri 40.8. In these flows, Ri is approximately constant with downslope distance, and decreases with increasing slope angle. Experiments show that U, and hence we, are also constant, so that the downslope flux Q and the lateral spreading increase linearly with distance. Downslope Flows into Stratified Environments
If the environmental fluid is density stratified, the effects of this stratification on density currents flowing over a horizontal surface are mostly limited to its effects on the surrounding flow, including the generation of internal waves (see Internal Waves). However, for flows down slopes, if ambient stratification is present it is a major parameter. If N is the
Figure 4 An example of flow of a layer of dense fluid down a slope at 61 to the horizontal, in a density-stratified environment, showing the structure of the interface. The dense fluid is the lightly colored dyed layer close to the boundary. Note the filaments extending from it, and the mixed fluid above. CM ¼ 0.015, with current depth of order 1 cm, buoyancy frequency of stratification N ¼ 1.24 rad s1, Q ¼ 2.9 cm2 s1, g0 ¼ 19.4 cm s2.)
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entrainment coefficient that depends on Ri, in a manner similar to eqn [10]. There is also a loss of dense fluid to the mixed layer, and an exchange between the mixed layer and the environment. Fluid may leave the mixed layer to find its own neutral level, and this occurs continuously along the path of the current. Hence dense fluid from the source may be distributed over a range of depths, and not all at one level. This detrainment into the environment depends on both Ri and M – larger M implies larger detrainment, to the extent that all the dense fluid may be detrained before it reaches its ambient level. The above sections describe the ‘pure’ dynamical form of a gravity current, in various different situations. But when gravity currents occur in the ocean, they are often strongly influenced by other factors over most of their life cycles. In particular, the flow of dense fluid down continental slopes is usually affected by the Earth’s rotation, which means that the flow is at least partly in geostrophic balance, with an Ekman layer next to the bottom boundary (see Rotating Gravity Currents). This applies in particular to the deep flows through Denmark Strait and around Antarctica, that drive the overturning thermohaline circulation. Also, sediment suspension is a common cause of submarine gravity currents, and it may also be a factor in these deep flows.
See also Internal Waves. Rotating Gravity Currents. Shelf Sea and Shelf Slope Fronts. Upper Ocean Mixing Processes. Upper Ocean Vertical Structure.
Further Reading Baines PG (1995) Topographic Effects in Stratified Flows. Cambridge: Cambridge University Press. Baines PG and Condie S (1998) Observations and modeling of Antarctic downslope flows: a review. In: Jacobs SS and Weiss R (eds.) Ocean, Ice and Atmosphere: Interactions at the Antarctic Continental Margin, AGU Antarctic Research Series, vol. 75, pp. 29--49. Washington, DC: American Geophysical Union. Simpson JE (1997) Gravity Currents. Cambridge: Cambridge University Press. Turner JS (1986) Turbulent entrainment: the development of the entrainment assumption, and its application to geophysical flows. Journal of Fluid Mechanics 173: 431--471.
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NORTH ATLANTIC OSCILLATION (NAO) J. W. Hurrell, National Center for Atmospheric Research, Boulder, CO, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1904–1911, & 2001, Elsevier Ltd.
Introduction Simultaneous variations in weather and climate over widely separated points on Earth have long been noted in the meteorological literature. Such variations are commonly referred to as ‘teleconnections’. In the extratropics, teleconnections link neighboring regions mainly through the transient behavior of atmospheric planetary-scale waves. Consequently, some regions may be cooler than average, while thousands of kilometers away warmer conditions prevail. Though the precise nature and shape of these structures vary to some extent according to the statistical methodology and the data set employed in the analysis, consistent regional characteristics that identify the most conspicuous patterns emerge. Over the middle and high latitudes of the northern hemisphere, a dozen or so distinct teleconnection patterns can be identified during boreal winter. One of the most prominent is the North Atlantic Oscillation (NAO). The NAO dictates climate variability from the eastern seaboard of the USA to Siberia and from the Arctic to the subtropical Atlantic. This widespread influence indicates that the NAO is more than just a North Atlantic phenomenon. In fact, it has been suggested that the NAO is the regional manifestation of a larger scale (hemispheric) mode of variability known as the Arctic Oscillation (see below). Regardless of terminology, meteorologists for more than two centuries have noted the pronounced influence of the NAO on the climate of the Atlantic basin. Variations in the NAO are important to society and the environment. Through its control over regional temperature and precipitation variability, the NAO directly impacts agricultural yields, water management activities, and fish inventories among other things. The NAO accounts for much of the interannual and longer-term variability evident in northern hemisphere surface temperature, which has exhibited a warming trend over the past several decades to values that are perhaps unprecedented over the past 1000 years. Understanding the processes that govern variability of the NAO is therefore of high priority,
especially in the context of global climate change. This article defines the NAO and describes its relationship to variations in surface temperature and precipitation, as well as its impact on variability in the North Atlantic Ocean and on the regional ecology. It concludes with a discussion of the mechanisms that might influence the amplitude and timescales of the NAO, including the possible roles of the stratosphere and the ocean.
What is the North Atlantic Oscillation? Like all atmospheric teleconnection patterns, the NAO is most clearly identified when time averaged data (monthly or seasonal) are examined, since time averaging reduces the ‘noise’ of small-scale and transient meteorological phenomena not related to large-scale climate variability. Its spatial signature and temporal variability are most often defined through the regional sea level pressure field, for which some of the longest instrumental records exist. The NAO refers to a north–south oscillation in atmospheric mass with centers of action near Iceland and over the subtropical Atlantic from the Azores across the Iberian Peninsula. Although it is the only teleconnection pattern evident throughout the year in the northern hemisphere, its amplitude is largest during boreal winter when the atmosphere is dynamically the most active. During the months December through March, for instance, the NAO accounts for more than one-third of the total variance in sea level pressure over the North Atlantic. A time series (or index) of more than 100 years of wintertime NAO variability and the spatial signature of the oscillation are shown in Figures 1 and 21 Differences of 415 hPa occur across the North Atlantic between the two phases of the NAO. In the socalled positive phase, higher than normal surface pressures south of 551N combine with a broad region of anomalously low pressure throughout the Arctic. Because air flows counterclockwise around low pressure and clockwise around high pressure in the northern hemisphere, this phase of the oscillation is associated with stronger than average westerly winds across the middle latitudes of the Atlantic onto
1 More sophisticated and objective statistical techniques, such as eigenvector analysis, yield time series and spatial patterns of average winter sea level pressure variability very similar to those shown in Figures 1 and 2.
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6
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0 _2 _4 _6 1870 1880 1890 1900 1910 1920 1930 1940 1950 1960 1970 1980 1990 2000
Figure 1 Winter (December–March) index of the NAO based on the difference of normalized sea level pressure between Lisbon, Portugal, and Stykkisholmur/Reykjavik, Iceland from 1864 to 2000. The average winter sea level pressure data at each station were normalized by division of each seasonal pressure by the long-term mean (1864–1983) standard deviation. The heavy solid line represents the index smoothed to remove fluctuations with periods o4 years.
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Figure 2 Difference in sea level pressure between years with an NAO index value 41.0 and those with an index value o minus low index winters) since 1899. The contour increment is 2 hPa and negative values are dashed.
Europe, with anomalous southerly flow over the eastern USA and anomalous northerly flow across western Greenland, the Canadian Arctic, and the Mediterranean.
1.0 (high
The NAO is also readily apparent in meteorological data throughout the depth of the troposphere, and its variability is significantly correlated with changes in the strength of the winter polar vortex in
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the stratosphere of the northern hemisphere. Within the lower stratosphere, the leading pattern of geopotential height variability is also characterized by a seesaw in mass between the polar cap and the middle latitudes, but with a much more zonally symmetric (or annular) structure than in the troposphere. When heights over the polar region are lower than normal, heights at nearly all longitudes in middle latitudes are higher than normal. In this phase, the stratospheric westerly winds that encircle the pole are enhanced and the polar vortex is ‘strong’ and anomalously cold. It is this annular mode of variability that has been termed the Arctic Oscillation. However, the signature of the stratospheric Arctic Oscillation in winter sea level pressure data looks very much like the anomalies associated with the NAO, with centers of action over the Arctic and the Atlantic (Figure 2). The ‘annular’ character of the Arctic Oscillation in the troposphere, therefore, reflects the vertically coherent fluctuations throughout the Arctic more than any coordinated behavior in the middle latitudes outside of the Atlantic basin. That the NAO and Arctic Oscillation reflect essentially the same mode of tropospheric variability is emphasized by the fact that their time series are nearly identical, with differences depending mostly on the details of the analysis procedure. There is little evidence for the NAO to vary on any preferred timescale (Figure 1). Large changes can occur from one winter to the next, and there is also a considerable amount of variability within a given winter season. This is consistent with the notion that much of the atmospheric circulation variability in the form of the NAO arises from processes internal to the atmosphere, in which various scales of motion interact with one another to produce random (and thus unpredictable) variations. On the other hand, there are also periods when anomalous NAO-like circulation patterns persist over many consecutive winters. In the Icelandic region, for instance, sea level pressure tended to be anomalously low during winter from the turn of the century until about 1930 (positive NAO index), while the 1960s were characterized by unusually high surface pressure and severe winters from Greenland across northern Europe (negative NAO index). A sharp reversal has occurred over the past 30 years, with strongly positive NAO index values since 1980 and sea level pressure anomalies across the North Atlantic and Arctic that resemble those in Figure 2. In fact, the magnitude of the recent upward trend is unprecedented in the observational record and, based on reconstructions using paleoclimate and model data, perhaps over the past several centuries as well. Whether such low frequency (interdecadal) NAO variability arises from
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interactions of the North Atlantic atmosphere with other, more slowly varying components of the climate system (such as the ocean), whether the recent upward trend reflects a human influence on climate, or whether the longer timescale variations in the relatively short instrumental record simply reflect finite sampling of a purely random process, are topics of considerable current interest.
Impacts of the North Atlantic Oscillation Temperature
The NAO exerts a dominant influence on wintertime temperatures across much of the northern hemisphere. Surface air temperature and sea surface temperature (SST) across wide regions of the North Atlantic Ocean, North America, the Arctic, Eurasia, and the Mediterranean are significantly correlated with NAO variability.2 Such changes in surface temperature (and related changes in rainfall and storminess) can have significant impacts on a wide range of human activities, as well as on marine and terrestrial ecosystems. When the NAO index is positive, enhanced westerly flow across the North Atlantic during winter moves relatively warm (and moist) maritime air over much of Europe and far downstream across Asia, while stronger northerlies over Greenland and north-eastern Canada carry cold air southward and decrease land temperatures and SST over the northwest Atlantic (Figure 3). Temperature variations over North Africa and the Middle East (cooling), as well as North America (warming), associated with the stronger clockwise flow around the subtropical Atlantic high-pressure center are also notable. The pattern of temperature change associated with the NAO is important. Because the heat storage capacity of the ocean is much greater than that of land, changes in continental surface temperatures are much larger than those over the oceans, so they tend to dominate average northern hemisphere (and global) temperature variability. Especially given the large and coherent NAO signal across the Eurasian continent from the Atlantic to the Pacific (Figure 3), it is not surprising that NAO variability explains about one-third of the northern hemisphere interannual surface temperature variance during winter. The strength of the link between the NAO and northern hemisphere temperature variability has also
2
Sea surface temperatures (SSTs) are used to monitor surface air temperature over the oceans because intermittent sampling is a major problem and SSTs have much greater persistence.
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Figure 3 Changes in land surface and sea surface temperatures ( 10 11C) corresponding to unit deviations of the NAO index for the winter months (December–March) from 1935 to 1999. The contour increment is 0.21C. Temperature changes 40.21C are indicated by dark shading, and those o 0.21C are indicated by light shading. Regions with insufficient data are not contoured.
added to the debate over our ability to detect and distinguish between natural and anthropogenic climate change. Since the early 1980s, winter temperatures over much of North America and Eurasia have been considerably warmer than average, while temperatures over the northern oceans have been slightly colder than average. This pattern, which has contributed substantially to the well-documented warming trend in northern hemisphere and global temperatures over recent decades, is quite similar to winter surface temperature changes projected by computer models forced with increasing concentrations of greenhouse gases and aerosols. Yet, it is clear from the above discussion that a significant fraction of this temperature change signal is associated with the recent upward trend in the NAO (Figures 1 and 3). How anthropogenic climate change might influence modes of natural climate variability such as the NAO, and the nature of the relationships between increased radiative forcing and interdecadal variability of these modes, remain central research questions. Precipitation and Storms
Changes in the mean circulation patterns over the North Atlantic are accompanied by changes in the
intensity and number of storms, their paths, and their associated weather. Here, the term ‘storms’ refers to atmospheric disturbances operating on timescales of about a week or less. During winter, a well-defined storm track connects the North Pacific and North Atlantic basins, with maximum storm activity over the oceans. The details of changes in storminess differ depending on the analysis method and whether the focus is on surface or upper-air features. Generally, however, positive NAO index winters are associated with a northward shift in the Atlantic storm activity, with enhanced activity from southern Greenland across Iceland into northern Europe and a modest decrease in activity to the south. The latter is most noticeable from the Azores across the Iberian Peninsula and the Mediterranean. Positive NAO winters are also typified by more intense and frequent storms in the vicinity of Iceland and the Norwegian Sea. The ocean integrates the effects of storms in the form of surface waves, so that it exhibits a marked response to long-lasting shifts in the storm climate. The recent upward trend toward more positive NAO index winters has been associated with increased wave heights over the north-east Atlantic and decreased wave heights south of 401N. Such
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changes have consequences for the operation and safety of shipping, offshore industries, and coastal development. Changes in the mean flow and storminess associated with swings in the NAO are also reflected in pronounced changes in the transport and convergence of atmospheric moisture and, thus, the distribution of evaporation and precipitation. Evaporation (E) exceeds precipitation (P) over much of Greenland and the Canadian Arctic during high NAO index winters (Figure 4), where changes between high and low NAO index states are on the order of 1 mm d 1. Drier conditions of the same magnitude also occur over much of central and southern Europe, the Mediterranean, and parts of the Middle East, whereas more precipitation than normal falls from Iceland through Scandinavia. This pattern, together with the upward trend in the NAO index since the late 1960s (Figure 1), is consistent with recent observed changes in precipitation over much of the Atlantic basin. One of the few regions of the world where glaciers have not exhibited a retreat over the past several decades is in Scandinavia, where more than average amounts of precipitation have been typical of many winters since the early 1980s. In contrast, over the Alps, snow depth and duration in recent winters have been among the lowest recorded this century, and the retreat of Alpine glaciers has been widespread. Severe drought has persisted throughout parts of Spain and Portugal as well. As far east as Turkey, river runoff is significantly correlated with NAO variability. There is also some observational and modeling evidence
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of a declining precipitation rate over much of the Greenland Ice Sheet over the past two decades, although measurement uncertainties are large. Ocean Variability
It has long been recognized that fluctuations in SST and the strength of the NAO are related, and there are clear indications that the North Atlantic Ocean varies significantly with the overlying atmosphere. The leading pattern of SST variability during winter consists of a tripolar structure marked, in one phase, by a cold anomaly in the subpolar North Atlantic, a warm anomaly in the middle latitudes centered off Cape Hatteras, and a cold subtropical anomaly between the equator and 301N. This structure suggests that the SST anomalies are driven by changes in the surface wind and air–sea heat exchanges associated with NAO variations. The relationship is strongest when the NAO index leads an index of the SST variability by several weeks, which highlights the well-known result that extratropical SST responds to atmospheric forcing on monthly and seasonal timescales. Over longer periods, persistent SST anomalies also appear to be related to persistent anomalous patterns of SLP (including the NAO), although the mechanisms which produce SST changes on decadal and longer time-scales remain unclear. Such fluctuations could primarily be the local oceanic response to atmospheric decadal variability. On the other hand, nonlocal dynamical processes in the ocean could also be contributing to the SST variations.
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Figure 4 Difference in evaporation (E) minus precipitation (P) between years with an NAO index value 41.0 and those with an index value o 1.0 (high minus low index winters) since 1958. The contour increment is 0.3 mm d 1. Differences 40.3 mm d 1 are indicated by dark shading, and those o0.3 mm d 1 are indicated by light shading.
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Subsurface ocean observations more clearly depict long-term climate variability, because the effect of the annual cycle and month-to-month variability in the atmospheric circulation decays rapidly with depth. These measurements are much more limited than surface observations, but over the North Atlantic they too indicate fluctuations that are coherent with the winter NAO index to depths of 400 m. The ocean’s response to NAO variability also appears to be evident in changes in the distribution and intensity of winter convective activity in the North Atlantic. The convective renewal of intermediate and deep waters in the Labrador Sea and the Greenland/ Icelandorwegian Seas contribute significantly to the production and export of North Atlantic Deep Water and, thus, help to drive the global thermohaline circulation. The intensity of winter convection at these sites is not only characterized by large interannual variability, but also interdecadal variations that appear to be synchronized with variations in the NAO (Figure 1). Deep convection over the Labrador Sea, for instance, was at its weakest and shallowest in the postwar instrumental record during the late 1960s. Since then, Labrador Sea Water has become progressively colder and fresher, with intense convective activity to unprecedented ocean depths (42300 m) in the early 1990s. In contrast, warmer and saltier deep waters in recent years are the result of suppressed convection in the Greeland/Icelandorwegian Seas, whereas intense convection was observed there during the late 1960s. There has also been considerable interest in the occurrence of low salinity anomalies that propagate around the subpolar gyre of the North Atlantic. The most famous example is the Great Salinity Anomaly (GSA). The GSA formed during the extreme negative phase of the NAO in the late 1960s, when clockwise flow around anomalously high pressure over Greenland fed record amounts of fresh water through the Denmark Strait into the subpolar North Atlantic ocean gyre. There have been other similar instances as well, and statistical analyses have revealed that these propagating salinity modes are closely connected to a pattern of atmospheric variability strongly resembling the NAO. Sea Ice
The strongest interannual variability of Arctic sea ice occurs in the North Atlantic sector. The sea ice fluctuations display an oscillation in ice extent between the Labrador and Greenland Seas. Strong interannual variability is evident in the sea ice
changes over the North Atlantic, as are longer-term fluctuations, including a trend over the past 30 years of diminishing (increasing) ice concentration during boreal winter east (west) of Greenland. Associated with the sea ice fluctuations are large-scale changes in sea level pressure that closely resemble the NAO. When the NAO is in its positive phase, the Labrador Sea ice boundary extends farther south, while the Greenland Sea ice boundary is north of its climatological extent. Given the implied surface wind changes (Figure 2), this is qualitatively consistent with the notion that sea ice anomalies are directly forced by the atmosphere, either dynamically via wind-driven ice drift anomalies, or thermodynamically through surface air temperature anomalies (Figure 3). The relationship between the NAO index (Figure 1) and an index of the North Atlantic ice variations is indeed strong, although it does not hold for all individual winters. This last point illustrates the importance of the regional atmospheric circulation in forcing the extent of sea ice. Ecology
Changes in the NAO have a wide range of effects on North Atlantic ecosystems. Temperature is one of the primary factors, along with food availability and spawning grounds, in determining the large-scale distribution pattern of fish and shellfish. Changes in SST and winds associated with changes in the NAO have been linked to variations in the production of zooplankton, as well as to fluctuations in several of the most important fish stocks across the North Atlantic. This includes not only longer-term changes associated with interdecadal NAO variability, but also interannual signals as well. Over land, fluctuations in the strength of the NAO have been linked to variations in plant phenology. In Norway, for example, most plant species have been blooming from 2–4 weeks earlier in recent years because of increasingly warm and wet winters, and many species have been blooming longer. Winter climate directly impacts the growth, reproduction, and demography of many animals, as well. European amphibians and birds have been breeding earlier over the past two to three decades, and these trends have been attributed to earlier growing seasons and increased forage availability. Variations in the NAO index are also significantly correlated with the growth, development, fertility, and demographic trends of large mammals from North America to northern Europe, such as northern ungulates and Canadian lynx.
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What are the Mechanisms that Govern influence on the Earth’s surface climate. Volcanic North Atlantic Oscillation Variability? aerosols act to enhance north–south temperature Although the NAO is an internal mode of variability of the atmosphere, surface, stratospheric or anthropogenic processes may influence its phase and amplitude. At present there is no consensus on the process or processes that most influence the NAO, especially on long (interdecadal) timescales. It is quite possible that NAO variability is affected by one or more of the mechanisms described below. Atmospheric Processes
Atmospheric general circulation model (AGCMs)3 provide strong evidence that the basic structure of the NAO results from the internal, nonlinear dynamics of the atmosphere. The observed spatial pattern and amplitude of the NAO are well simulated in AGCMs forced with fixed climatological annual cycles of solar insolation and SST, as well as fixed atmospheric trace gas composition. The governing dynamical mechanisms are interactions between the time-mean flow and the departures from that flow. Such intrinsic atmospheric variability exhibits little temporal coherence and, indeed, the timescales of observed NAO variability (Figure 1) do not differ significantly from this reference. A possible exception is the interdecadal NAO variability, especially the strong trend toward the positive index polarity of the oscillation over the past 30 years. This trend exhibits a high degree of statistical significance relative to the background interannual variability in the observed record; moreover, multi-century AGCM experiments like those described above do not reproduce interdecadal changes of comparable magnitude. A possible source of the trend could be through a connection to processes that have affected the strength of the atmospheric circulation in the lower stratosphere. During winters when the stratospheric westerlies are enhanced, the NAO tends to be in its positive phase. There is a considerable body of evidence to support the notion that variability in the troposphere can drive variability in the stratosphere, but it also appears that some stratospheric control of the troposphere may also occur. The atmospheric response to strong tropical volcanic eruptions provides some evidence for a stratospheric
3 Atmospheric general circulation models consist of a system of equations that describe the large-scale atmospheric balances of momentum, heat, and moisture, with schemes that approximate small-scale processes such as cloud formation, precipitation, and heat exchange with the sea surface and land.
gradients in the lower stratosphere by absorbing solar radiation in lower latitudes. In the troposphere, the aerosols exert only a very small direct influence. Yet, the observed response following eruptions is not only lower geopotential heights over the pole with stronger stratospheric westerlies, but also a positive NAO-like signal in the tropospheric circulation. Reductions in stratospheric ozone and increases in greenhouse gas concentrations also appear to enhance the meridional temperature gradient in the lower stratosphere, leading to a stronger polar vortex. It is possible, therefore, that the upward trend in the NAO index in recent decades (Figure 1) is associated with trends in either or both of these quantities. Indeed, a decline in the amount of ozone poleward of 401N has been observed during the last two decades, and the stratospheric polar vortex has become colder and stronger. Ocean Forcing of the Atmosphere
In the extratropics, it is clear that the atmospheric circulation is the dominant driver of upper ocean thermal anomalies. A long-standing issue, however, has been the extent to which anomalous extratropical SST feeds back to affect the atmosphere. Most evidence suggests that this effect is quite small compared with internal atmospheric variability. Nevertheless, the interaction between the ocean and atmosphere could be important for understanding the details of the observed amplitude of the NAO, its interdecadal variability, and the prospects for meaningful predictability. While intrinsic atmospheric variability is random in time, theoretical and modeling evidence suggest that the ocean can respond to it with marked persistence or even oscillatory behavior. On seasonal and interannual timescales, for example, the large heat capacity of the upper ocean can lead to slower changes in SST relative to the faster, stochastic atmospheric forcing. On longer timescales, SST observations display a myriad of variations. Middle and high latitude SSTs over the North Atlantic, for example, were colder than average during the 1970s and 1980s, but warmer than average from the 1930s through the 1950s. Within these periods, shorterterm variations in SST are also apparent, such as a dipole pattern that fluctuates on approximately decadal time scales with anomalies of one sign east of Newfoundland, and anomalies of opposite polarity off the south-east coast of the USA. A key to whether or not changes in the state of the NAO reflect these variations in the state of the ocean
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surface is the sensitivity of the atmosphere to middle and high latitude SST and upper ocean heat content anomalies. Most AGCM studies show weak responses to extratropical SST anomalies, with sometimes contradictory results. Yet, some AGCMs, when forced with the time history of observed, global SSTs and sea ice concentrations over the past 50 years or so, show modest skill in reproducing aspects of the observed NAO behavior, especially its interdecadal fluctuations (Figure 1). Such results do not necessarily imply, however, that the extratropical ocean is behaving in anything other than a passive manner. It could be, for instance, that long-term changes in tropical SSTs force a remote atmospheric response over the North Atlantic, which in turn drives changes in extratropical SSTs and sea ice. Some model studies indicate a sensitivity of the North Atlantic atmosphere to tropical SST variations, including variations over the tropical Atlantic which are substantial on both interannual and interdecadal timescales. The response of the extratropical North Atlantic atmosphere to changes in tropical and extratropical SST distributions, and the role of land processes and sea ice in producing atmospheric variability, are problems which are currently being addressed. Until these are better understood, it is difficult to evaluate the realism of more complicated scenarios that rely on truly coupled interactions between the atmosphere, ocean, land, and sea ice to produce North Atlantic climate variability. It is also difficult to evaluate the extent to which interannual and longerterm variations of the NAO might be predictable.
Glossary Great salinity anomaly A widespread freshening of the upper 500–800 m layer of the far northern North Atlantic Ocean, traceable around the subpolar gyre from its origins north of Iceland in the mid-to-late 1960s until it return to the Greenland Sea in the early 1980s.
Fisheries and Climate. Heat and Momentum Fluxes at the Sea Surface. North Sea Circulation. Ocean Circulation. Open Ocean Convection. Plankton and Climate. Sea Ice. Sea Ice: Overview. Wave Generation by Wind. Wind Driven Circulation.
Further Reading Appenzeller C, Stocker TF, and Anklin M (1998) North Atlantic oscillation dynamics recorded in Greenland ice cores. Science 282: 446--449. Dickson B (1999) All change in the Arctic. Nature 397: 389--391. Hurrell JW (1995) Decadal trends in the North Atlantic Oscillation regional temperatures and precipitation. Science 269: 676--679. Kerr RA (1997) A new driver for the Atlantic’s moods and Europe’s weather? Science 275: 754--755. National Research Council (1998) Decade-to-century scale climate variability and change. A science strategy. Washington: National Academy Press. Rodwell MJ, Rowell DP, and Folland CK (1999) Oceanic forcing of the wintertime North Atlantic Oscillation and European climate. Nature 398: 320--323. Shindell DT, Miller RL, Schmidt G, and Pandolfo L (1999) Simulation of recent northern winter climate trends by greenhouse gas forcing. Nature 399: 452--455. Stenseth NC and Co-authors 1999 (1999) Common dynamic structure of Canadian lynx populations within three climatic regions. Science 285: 1071--1073. Sutton R and Allen MR (1997) Decadal predictability of North Atlantic sea surface temperature and climate. Nature 388: 563--567. Thompson DWJ and Wallace JM (1998) The Arctic oscillation signature in the wintertime geopotential height and temperature fields. Geophysical Research Letters 25: 1297--1300. World Climate Research Programme (1998) The North Atlantic Oscillation. Climate Variability and Predictability (CLIVAR) Initial Implementation Plan. WCRP no. 103, WMO/TD no. 869, ICPO no. 14, 163–192.
See also Abrupt Climate Change. Arctic Ocean Circulation. Current Systems in the Atlantic Ocean. Elemental Distribution: Overview. Evaporation and Humidity.
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NORTH SEA CIRCULATION M. J. Howarth, Proudman Oceanographic Laboratory, Wirral, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 1, pp 1912–1921, & 2001, Elsevier Ltd.
Introduction Currents in the North Sea, as in any continental shelf sea, occur in response to forcing by tides, winds, density gradients (arising from freshwater input), and pressure gradients. For most of the North Sea the dominant motion is tidal, with the next most significant motion generated by wind forcing. The response is determined by the sea’s topography and bathymetry and is modified by, and modifies, its density distribution. Currents occur over a range of scales – a practical minimum is indicated in each case, ignoring waves and turbulence which are outside the scope of this article. For time the range is from minutes to years, horizontally in space from o1 km to basin-wide (1000 km), and vertically from 1 m to the full water depth. Current amplitudes range generally from 0.01 m s 1 to 2 m s 1 (in a very few places to in excess of 5 m s 1). The North Sea has been extensively studied for more than 100 years and its physical oceanography has been widely reviewed (see Further Reading section).
Topography The North Sea is a semi-enclosed, wide continental shelf sea (Figure 1). It stretches southward for about 1000 km from the Atlantic Ocean in the north, and is about 500 km broad, from west to east. The northern boundary is composed of two connections. Firstly, the main connection with the ocean at the shelf edge is between the Shetland Islands and Norway. Secondly, from the mainland of Scotland to the Orkney and Shetland Islands via the Pentland Firth and the Fair Isle Channel, the connection is with the continental shelf seas to the west and north of Scotland. The North Sea has two other, narrow, connections. One is at its southern end, via the Dover Strait to the relatively salty English Channel (and ultimately again Atlantic) water. The second, on its eastern side, is via the Skagerrak and Kattegat to fresh Baltic water. Topographically the North Sea can be split into three (Figure 2). In the south, bounded by about 541N and the shallow Dogger Bank, depths are o40 m.
Secondly, to the north of this the water depth increases from 40 m to 200 m at the shelf edge. Thirdly, the deep Norwegian Trench is in the north east, penetrating from the Norwegian Sea southward along the coast of Norway. The deepest depths in the trench, about 750 m, occur in the Skagerrak, off southern Norway and western Sweden, whilst the minimum axial depth is about 225 m off western Norway. This has a significant impact on the sea’s dynamics.
Meteorology The North Sea lies at temperate latitudes, between 511N and 621N, and so experiences pronounced seasonal changes in meteorological conditions. From September to April, in particular, it is exposed to the effect of a series of storms. These usually track eastward to the north of the British Isles, taking about a day to pass by the North Sea, accompanied by strong winds from the south west, west, and north west. Although these are the dominant wind directions, strong winds can blow from any direction. To some extent the British Isles shelter the North Sea from the wind’s full effect since the only wind direction uninterrupted by land is from the north. (This is more significant for waves than currents.) The size of the storms is generally of the same order as that of the North Sea, or larger, and they can recur every few days. On average gale force winds blow on 30 days per year. The extreme ‘one-in-50-year’ hourly mean wind speeds are estimated to decrease from 39 m s 1 in the north to 32 m s 1 in the south, slightly stronger in the center of the North Sea compared with near the coast.
River Discharges The main input of fresh water into the North Sea is from rivers, including the Rhine and Elbe, discharging along its southern coast (Belgium, The Netherlands, Germany). On average this input is 4000 m3 s 1, but it is highly variable from day to day and also from year to year. This fresh water has a significant impact on the salinity and density distribution (lower salinities along the continental coast and in the German Bight) and ultimately on the longterm circulation. A further 1500 m3 s 1 of fresh water on average is discharged from Scottish and English rivers. More fresh water flows into the Kattegat from the Baltic, on average 15 000 m3 s 1 in the form of a low salinity (8.7 PSU) flow of
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Depth, m 0_30 30_40 40_50 50_100 100_200 200_300 300_400 400_500 500_600 600_700 700_800 800_900 900_1000 1000_1100 1100_1200 1200_1300 1300_1400 Below 1400
Figure 2 Bathymetry of the North Sea. (Reproduced with permission from North Sea Task Force, 1993.)
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NORTH SEA CIRCULATION
30 000 m3 s 1 in a surface layer partially compensated by a flow of 15 000 m3 s 1 of higher salinity (17.4 PSU) North Sea water into the Baltic in a nearbed layer. However, the instantaneous exchange between the Kattegat and the Baltic Sea is by no means a steady two-layer flow, but is predominantly barotropic (in response to wind forcing and the resulting sea level slopes) and can be up to 300 000 m3 s 1 in either direction. The fresher water flows into the Norwegian Coastal Current, a surface current following the Norwegian coast, and exits the North Sea. Generally there is little exchange with water in the rest of the North Sea, although on occasion a surface layer of low salinity water can spread westward across much of the northern North Sea. The boundary between the Norwegian Coastal Current and Atlantic water in the northern North Sea (see Figure 3) is not smooth but is subject to meanders with a wavelength of 50–100 km which can break off into gyres and vortex pairs generally propagating northward. Orbital speeds within the gyres are asymmetric; up to 2 m s 1 has been reported in the upper layer.
Stratification The behavior of currents depends on whether the water column is well mixed or stratified. For some processes, for instance the response to wind forcing, stratification leads to the surface layer becoming decoupled from the bed layer – the sharper the transition from surface to bed layers (called the thermocline, for temperature) the greater the degree of decoupling. Well-mixed and stratified regions are separated by sharp fronts. Stratification arises when mixing is insufficient to mix down lighter water at the sea surface – the water might be lighter either because of solar heat input during summer or because of freshwater river discharge. The energy available for mixing is primarily derived from the tides and its effect on the water column depends on the water depth. Hence the southern North Sea, where depths are shallow and tidal currents strong, tends to remain well mixed throughout the year, apart from regions close to the coast affected by river plumes. Solar heat input causes most of the northern North Sea to stratify between April/May and October/ December, with a well-mixed surface layer of about 30–40 m deep. In autumn heat loss at the surface leads to the surface mixed layer deepening and cooling until the bottom is reached. Stratification caused by river discharge can occur at any time of the year, the fresher water tending to
form a thin surface layer in a plume about 30 km wide which stays close to the coast. The nature of the plume is very variable, depending on freshwater discharge, wind, and tide (both during the semi–diurnal and the spring-neap cycles). Mean currents in the plume can be up to 0.2 m s 1.
Measurements (see Drifters and Floats; Moorings; Nuclear Fuel Reprocessing and Related Discharges; Profiling Current Meters; Single Point Current Meters) The long-term circulation (Figure 3) has been investigated since before 1900 by studying the returns of surface and seabed drifters and the distribution of tracers, initially salinity. More recently dissolved radionuclides have been used as tracers, primarily cesium-137 and technetium-99 discharged into the sea at very low levels in waste from nuclear fuel reprocessing plants, mainly at Sellafield on the west coast of the UK and Cap de la Hague on the north coast of France. These have given unambiguous confirmation of circulation paths. Since the 1960s variations in currents on short time scales have been measured at fixed points by deploying moorings, often for about a month, with current meters, fitted with rotors or propellers, recording data every few minutes. In the 1990s a new instrument, acoustic Doppler current profiler (ADCP), became available which can measure the current profile throughout the majority of the water column in continental shelf seas. Frequently deployed in a seabed frame the current is determined from the Doppler shift in the back-scattered signal transmitted at 201 or 301 to the vertical. Currents at different depths are determined by chopping the return signal into segments. Another new technique is shore-based hf radar, where again the Doppler shift in a back-scattered signal is used to determine surface currents over a region, typically in 1 km cells, out to a range of about 30 km.
Numerical Models At any one time measurements can give only an estimate of currents with very limited spatial coverage compared with the extent of the North Sea. However, these can be complemented by numerical models to give fuller spatial coverage. The basic hydrodynamic equations are well known and have been solved on regular finite difference or irregular finite element grids. Currents vary in two horizontal directions and in the vertical, but the models are particularly quick if the equations are depth-averaged,
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which is especially useful for the estimation of tide and storm surge elevations. Such models are now run operationally in association with meteorological models (the only sufficiently detailed source of meteorological forcing data to drive realistic sea models). Three-dimensional models are being developed which predict the tidal and wind-driven current profile and the effects of horizontal and vertical density gradients. Given the difficulty and expense of making accurate current measurements the future for prediction is via numerical models tested against a few critical measurements. Numerical models are also an integral component of process studies, since the significance of terms in the equations can be isolated and determined.
Tides The dominant motion for most of the shelf seas around the UK is the semi-diurnal tides in response to the gravitational attraction of the moon and the sun. The tides are regular, repetitive, and predictable. Tidal currents control many physical and nonphysical processes and determine mixing, even though they are predominantly cyclical (to and fro or describing an ellipse) and tend to have small net movement, typically 0.01–0.03 m s 1, except near irregularities in the coastline such as headlands. In addition to the familiar basic M2 12.4 hour cycle, the tides experience fortnightly spring–neap (on average spring tidal currents are twice as big as neap tidal currents), monthly, and 6 monthly cycles (larger spring tides usually occur near the spring and autumn equinoxes). Away from coasts the tidal currents can be accurately approximated (better than 0.1 m s 1) by just four constituents, or frequencies – M2, S2, N2, and K2, with periods between 11.97 and 12.66 hours. Near the shelf edge there can be localized amplification of the diurnal constituents O1 and K1. Only close to the coasts, in shallow water and near headlands, and in estuaries are distortions to the tides (represented by higher harmonics) significant. Clearly this is a very important region for predicting sea levels for navigation, but represents a small proportion of the area of the North Sea. The tides enter the North Sea from the Atlantic Ocean north of Scotland and sweep round it in anticlockwise sense as a progressive (Kelvin) wave gradually losing energy by working against bottom friction. By the time Norway is reached tidal currents are relatively weak, o 0.1 m s 1 at spring tides. Maximum currents occur in localized coastal constrictions, such as the Pentland Firth (in excess of 5 m s 1 at spring tides) and the Dover Strait (2 m s 1). Elsewhere currents at spring tides rarely exceed 1 m s 1 (Figure 4).
The amplitude of tidal currents does not vary significantly with depth, except in a boundary layer near to the bed generated by friction, of the order of a few meters (up to 10 m) thick. Here the current profile at maximum flow is approximately logarithmic in distance from the bottom. Also due to the effects of bottom friction tidal currents near the bed are in advance of those near the surface, so that the tide turns earlier, by up to half an hour, near the bed compared with near the surface.
Wind Forcing Wind forcing is responsible for the second largest currents after tides over most of the region, and in areas where tidal currents are weak (for instance off Norway), the largest. By the very nature of storms the forcing is intermittent and seasonal. The largest storms tend to occur in winter, when the water column is homogeneous. Two effects are important. Firstly, sea level gradients are set up against a coast counterbalancing the wind stress, since the North Sea is semi-enclosed and so has no long approximately straight coasts, in contrast with many other continental shelf seas. Secondly there is the direct action of the wind stress at the sea surface. The sea level gradients lead to storm surges at the coast, with the threat of coastal flooding. Sizeable currents can be generated during the setting up and relaxing of the sea level gradients. Measurement of these currents at depth is scarce not only because the environment is harsh during storms, although less severe than at the surface, but also because long deployments are required – at least over a 6 month period including winter to ensure an adequate sampling of storms. Seabed-mounted ADCPs are helpful, both because the storm environment is less likely to impinge on data quality and also because longer measurements are possible. The only feasible way of predicting extreme currents, for instance the ‘one-in50-year’ current for the design of offshore structures, is via numerical models, either run for long periods or forced by a series of extreme events. The accuracy of extreme sea level estimates is better than that for extreme currents, because models can be validated against long records (many years) of coastal elevations from tide gauges. The direct current response at the sea surface to the action of the wind stress is very difficult to measure, since waves tend to be large when winds are strong, but it appears to be limited to at most the top 25 m of the water column. (A very useful remote sensing technique here is hf radar which measures surface currents.) The response is rapid, less than a
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60°N
55°N
5°N
°W
0°
5°E
Figure 4 Estimate of the maximum depth-averaged currents for an average spring tide, in m s 1.
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10°E
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NORTH SEA CIRCULATION
few hours, and dependent on the roughness of the sea surface (i.e. on the waves and whitecapping). The surface current amplitude used, for instance in oil spill studies, is usually taken to be in the range 1–6% of the wind speed, with a typical value of 3%. The higher ratio appears to be more appropriate for developing seas. The direction of the wind drift is a few degrees to the right of the wind. Since there are very few accurate measurements to test models critically, the accuracy of models in this region is uncertain.
Circulation Circulation is the long-term movement of water. Within the North Sea it is forced by the net tides (for instance, responsible for 50% of water transport in the western North Sea), the mean winds, and density gradients, principally caused by freshwater input from rivers and the Baltic. All these forces tend to act in the same sense, described below. There are seasonal variations, as storms mainly occur in winter, river discharge has an annual cycle, and in summer the water column stratifies in some regions. Over short periods the net movement of water is likely to be determined by the last storm or storms. The circulation pattern outlined below is enhanced by winds from the south west and can be reversed by strong winds from south and east. Although of considerable and long-standing interest to oceanographers – to physicists, fisheries biologists, and to those studying the movement of suspended sediments and pollution – it is difficult to measure and quantify since in most places its magnitude is o 0.1 m s 1, much smaller than tidal or extreme wind-driven currents. It frequently has very short space scales, changing significantly in magnitude over a few kilometers and changing direction by 1801 from the surface to the seabed. There is, however, an overall long-term mean pattern (Figure 3), shown up clearly in the distributions of tracers whose major features in terms of direction have been known since the early 1900s. There are large uncertainties on estimates of amplitudes. The flow is broadly anticlockwise round the coasts of the North Sea, with weak and varied circulation in its center. The mean coastal flow is southward past Scotland and England into the Southern Bight, where there are inputs of salty water through the Dover Straits and of fresh water down the main rivers, all of which pass through large industrialized regions (Humber, Thames, Meuse, Scheldt, Rhine, Weser, Elbe) and on into the German Bight, flowing northward past Denmark in the Jutland current to join the Norwegian Coastal Current
in the Skagerrak. The average input through the Dover Strait is about 0.1 106 m3 s 1, but the flow there is strongly influenced by the wind and can reverse, moving south-westward out of the North Sea. There is also some flow from the coast of East Anglia to the north Dutch coast. There are major inflows (in excess of 106 m3 s 1) of water of Atlantic origin across the northern boundary, but very little penetrates far into the North Sea. The larger portion flows in on the western side of the Norwegian Trench and recirculates in the Skagerrak, flowing out along the eastern side of the Norwegian Trench underneath the Norwegian Coastal Current. A smaller inflow of mixed Atlantic and shelf water flows in on either side of the Shetland Islands and turns eastward to follow the 100 m contour across the North Sea, eventually to flow out along the eastern side of the Norwegian trench. The majority of water passing through the North Sea goes through the Skagerrak. The only major outflow is along the eastern side of the Norwegian Trench, with magnitude of about (1.3–1.8) 106 m3 s 1.
Effects of Stratification Stratification profoundly affects the nature of currents in two ways. Firstly currents are driven by the density gradients associated with fronts – the sharp boundaries between stratified and well-mixed water. The currents tend to be along the fronts with transverse exchanges inhibited. Secondly, the structure of motion in the water column is affected by the region of vertical density change – either because the surface and bed layers are decoupled or because internal waves can propagate along the density interface. Theoretically there is no difference between stratification caused by heat input and that caused by fresh water; dynamically all that matters is the density difference. In practice the consequences of thermal stratification in summer can be easier to study because the extent of the surface mixed layer and of the water depth are more manageable and because regions of thermal stratification are predictable and dependable. Fronts (see Shelf Sea and Shelf Slope Fronts)
The main front separating the thermally stratified water in summer to the north from the well-mixed water to the south starts from Flamborough Head, bifurcates around the Dogger Bank and passes to the north of the Frisian Islands. Tidal current speeds and water depths determine the front’s position. Currents along the front in a west to east sense are typically about 0.05 m s 1 but can be up to 0.15 m s 1.
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Inertial Currents
Shelf Edge
Inertial currents are generated by pulses of wind. The currents rotate clockwise in the Northern Hemisphere with a period in hours of 12/sine (latitude), which, in the North sea, is from 13.6 h at 621N to 14.6 h at 551N. In stratified water their vertical structure is primarily the first baroclinic mode; the currents are 1801 out of phase above and below the thermocline and integrate to zero over the water column. Because of the phase reversal large shears can be generated across the thermocline – the largest in the water column. Their decay time, if not reinvigorated by another wind pulse, is of the order of days. ADCPs covering the majority of the water column are ideal instruments for measuring inertial currents, which can show all the characteristics indicated above. In the North Sea inertial current speeds up to 0.25 m s 1 have been observed in the surface well-mixed layer, which for most of the period of stratification is the shallower and hence the layer where inertial currents are stronger. During stratification some energy at or near inertial periods is present most of the time, but not always with a simple first baroclinic mode structure. The situation is different from the open ocean – in shelf seas the water depth is limited and tidal friction/mixing is ever present.
The main distinguishing feature of the shelf edge is the pronounced change in water depth over a short distance, from shelf depths (o 200 m) to oceanic (usually several thousand meters). At the shelf break between the North Sea and the Faroe–Shetland Channel water depths deepen from 200 m to 1000– 1500 m over a distance of tens of kilometers. Compared with other shelf edges, even around the UK, this is relatively gentle and the gradients are relatively smooth. A mixture of dynamical processes can exist at the shelf edge involving a combination of stratification and steep bottom slope. Three relevant to the North Sea are internal tides (see above), the slope current flowing north-westward, and coastal trapped waves with super-inertial periods which can, for instance, lead to enhanced diurnal tides (see above).
Internal waves (see Internal Waves)
Internal waves are ubiquitous in stratified water. They propagate along the thermocline density interface as a cyclical vertical movement of the density interface, which can be quite large, over 10 m, with only a 0 (1 cm) movement of the sea surface. The restoring force is primarily the buoyancy force. Internal interfacial waves are analogous to surface waves but their phase speed is less. They can have any period between the local inertial (see above) and the Brunt-Va¨issa¨la¨ or buoyancy period, which depends on the degree of stratification and can be as short as a few minutes for sharp stratification. The corresponding currents are out of phase above and below the density interface and, for small amplitude waves, sum to zero in the vertical. It is only possible to give generalities since each situation is unique. Because vertical accelerations are significant, nonhydrostatic numerical models are required to reproduce the dynamics. These are rare for shelf-wide models. Internal tides are often generated as the tides cross sharp changes in bathymetry, such as the shelf break. The internal tide then propagates both into the shelf sea and into the ocean, usually in the form of a pulse of waves with periods of order 30 minutes, repeated with a tidal periodicity.
Conclusions The North Sea is a typical semi-enclosed continental shelf sea, with a wide range of tidal conditions, winds, circulation, and effects of stratification. To interpret and understand the dynamics it is not sufficient just to study the currents but also seabed pressures/coastal elevations, the (horizontal and vertical) density field which is not static, meteorological forcing (wind, atmospheric pressure, heat input at sea surface), and river discharges.
See also Drifters and Floats. Moorings. Nuclear Fuel Reprocessing and Related Discharges. Regional and Shelf Sea Models. Shelf Sea and Shelf Slope Fronts. Single Point Current Meters. Storm Surges. Tides.
Further Reading Charnock H, Dyer KR, and Huthnance JM (eds.) (1994) Understanding the North Sea System. London: Chapman and Hall. Eisma D (1987) The North Sea: an overview. Philosophical Transactions of the Royal Society B 316: 461--485. Lee AJ (1980) North Seap: hysical oceanography. Elsevier Oceanography Series 24B: 467--493. North Sea Task Force (1993) North Sea Quality Status Report 1993. Oslo and Paris Commissions, London. Fredensborg, Denmark: Olsen Olsen. Otto L, Zimmerman JTF, and Furnes GK (1990) Review of the physical oceanography of the North Sea. Netherlands Journal of Sea Research 26: 161--238. Rodhe J (1998) The Baltic and North Seasa: processoriented review of the physical oceanography. In: Robinson AR and Brink KH (eds.) The Sea, 11, pp. 699--732. Chichester: John Wiley Sons.
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NUCLEAR FUEL REPROCESSING AND RELATED DISCHARGES H. N. Edmonds, University of Texas at Austin, Port Aransas, TX, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1921–1928, & 2001, Elsevier Ltd.
Introduction Oceanographers have for some decades made use of human perturbations to the environment as tracers of ocean circulation and of the behavior of similar substances in the oceans. The study of the releases of anthropogenic radionuclides – by-products of the nuclear industry – by nuclear fuel reprocessing plants is an example of such an application. Other examples of the use of anthropogenic substances as oceanographic tracers addressed in this Encyclopedia include tritium and radiocarbon, largely resulting from atmospheric nuclear weapons testing in the 1950s and 1960s, and chlorofluorocarbons (CFCs), a family of gases which have been used in a variety of applications such as refrigeration and polymer manufacture since the 1920s. As outlined in this article, the source function of nuclear fuel reprocessing discharges, which is very different from that of either weapons test fallout or CFCs, is both a blessing and a curse, defining the unique applicability of these tracers but also requiring much additional and careful quantification. The two nuclear fuel reprocessing plants of most interest to oceanographers are located at Sellafield, in the UK, which has been discharging wastes into the Irish Sea since 1952, and at Cap de la Hague, in north-western France, which has been discharging wastes into the English Channel since 1966. These releases have been well documented and monitored in most cases. It should be noted, however, that the releases of some isotopes of interest to oceanographers have not been as well monitored throughout the history of the plants because they were not of radiological concern, were difficult to measure, or both. This is particularly true for 99Tc and 129I, which are the isotopes currently of greatest interest to the oceanographic community but for which specific, official release data are only available for the last two decades or less. The radionuclides most widely studied by ocean scientists have been 137 Cs, 134Cs, 90Sr, 125Sb, 99Tc, 129I, and Pu isotopes.
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The Sellafield plant has dominated the releases of most of these, with the exception of 125Sb and, particularly in the 1990s, 129I. In general, the Sellafield releases have been particularly well documented and are the easiest for interested scientists and other parties to obtain. The figures of releases presented in this article are based primarily on the Sellafield data, except where sufficient data are available for Cap de la Hague. The location of the Sellafield and Cap de la Hague plants is important to the oceanographic application of their releases. As discussed below, the releases enter the surface circulation of the high latitude North Atlantic and are transported northwards into the Norwegian and Greenland Seas and the Arctic Ocean. They have been very useful as tracers of ventilation and deep water formation in these regions, which are of great importance to global thermohaline circulation and climate. In addition, the study of the dispersion of radionuclides from the reprocessing plants at Sellafield and Cap de la Hague serves as an analog for understanding the ultimate distribution and fate of other wastes arising from industrial activities in north-western Europe. Much of the literature on reprocessing releases in the oceans has focused on, or been driven by, concerns relating to contamination and radiological effects, and they have not found the same broad or general oceanographic application as some other anthropogenic tracers despite their great utility and the excellent quality of published work. This may be attributable to a variety of factors, including perhaps (1) the contaminant-based focus of much of the work and its publication outside the mainstream oceanographic literature, (2) the complications of the source function as compared with some other tracers, and (3) the comparative difficulty of measuring many of these tracers and the large volumes of water that have typically been required. This article addresses the historic and future oceanographic applications of these releases. A Note on Units
In most of the literature concerning nuclear fuel reprocessing releases, amounts (both releases and ensuing environmental concentrations) are expressed in activity units. The SI unit for activity is the becquerel (Bq), which replaced the previously common curie
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NUCLEAR FUEL REPROCESSING AND RELATED DISCHARGES
(Ci), a unit which was based on the activity of one gram of radium. These two units are related as follows: 1Bq ¼ 1disintegration per second 1Ci ¼ 3:7 1010 Bq The conversion from activity to mass or molar units is dependent on the half-life of the isotope in question. Activity (A) is the product of the number of atoms of the isotope present (N) and its decay constant (l): A ¼ lN. The decay constant, l, is related to the half-life by t1=2 ¼ ðln2Þ=l. Using this equation, it is simple to convert from Bq to mol, given the decay constants in units of s1, and Avogadro’s number of 6.023 1023 atoms mol1. It is also worth noting that the total activity of naturally occurring radionuclides in sea water is approximately 12 kBq m3. Most of this activity is from the long-lived naturally occuring isotope, 40 K(t1/2 ¼ 1.25 109 years). Activities contributed from anthropogenic radionuclides greatly exceed this (millions of Bq m3) in the immediate vicinity of the Sellafield and La Hague outfall pipes. However, most oceanographic studies of these releases involve measurements of much lower activities: up to 10s of Bq m3 in the case of 137Cs, and generally even lower for other radionuclides, for instance mBq m3 or less for plutonium isotopes and 129I.
Nuclear Fuel Reprocessing and Resulting Tracer Releases Origin and Description of Reprocessing Tracers
Nuclear fuel reprocessing involves the recovery of fissile material (plutonium and enriched uranium) and the separation of waste products from ‘spent’ (used) fuel rods from nuclear reactors. In the process, fuel rods, which have been stored for a time to allow short-lived radionuclides to decay, are dissolved and the resulting solution is chemically purified and separated into wastes of different composition and activity. Routine releases from the plants to the environment occur under controlled conditions and are limited to discharge totals dictated by the overseeing authorities of each country. The discharges have varied over the years as a function of the amount and type of fuel processed and changes in the reprocessing technology. The discharge limits themselves have changed, in response to monitoring efforts and also spurred by technological advances in waste-treatment capabilities. As a general rule, these changes have resulted in decreasing releases for most nuclides (most notably cesium and the actinide elements), but
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there are exceptions. For instance, the Enhanced Actinide Removal Plant (EARP) was constructed at the Sellafield site in the 1990s in order to enable the additional treatment and subsequent discharge of a backlog of previously stored wastes. Although the new technology enabled the removal of actinide elements from these wastes, it is not effective at removing 99Tc. Therefore an allowance was made for increased discharge of 99Tc, up to 200 TBq per year. The resulting pulse of increased 99Tc discharges from Sellafield beginning in 1994 is currently being followed with great interest, as is discussed in more detail below. The end result of these processes is the availability of a suite of oceanographic tracers with different discharge histories (e.g., in terms of the timing and magnitude of spikes) and a range of half-lives, and thus with a range of utility and applicability to studies of oceanographic processes at a variety of spatial and temporal scales. A brief summary of the reprocessing radionuclides most widely applied in oceanography is shown in Table 1, and examples of discharge histories are given in Figure 1. It has also been particularly useful in some cases to measure the ratio of a pair of tracers, for instance 134Cs/137Cs, 137 Cs/90Sr, or potentially, 99Tc/129I. The use of isotope ratios can (1) provide temporal information and aid the estimation of rates of circulation, (2) mitigate the effects of mixing which will often alter individual concentrations more than ratios, and (3) aid in distinguishing the relative contributions of different sources of the nuclides in question. Finally, differences in the chemical behavior of the different elements released can be exploited to study different processes. Although most of the widely applied tracers are largely conservative in sea water and therefore serve as tracers of water movement, the actinides, particularly Pu, have a high affinity for particulate material and accumulate in the sediments. These tracers are then useful for studying sedimentary processes. The Reprocessing Tracer Source Function
The primary difference between the north-western European reprocessing releases and other anthropogenic tracers used in oceanography is the nature of their introduction to the oceans. Reprocessing releases enter the oceans essentially at a point source, rather than in a more globally distributed fashion as is the case for weapons test fallout or the chlorofluorocarbons. Thus, reprocessing tracers are excellent, specific tracers for waters originating from north-western Europe. Because these waters are transported to the north into the Nordic Seas and
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Summary of the major reprocessing tracers and their applications
Table 1 Isotope
Half-life (years)
Sources
Applications
137
30
Weapons testing, reprocessing (mostly Sellafield), Chernobyl
90
28
Weapons testing, reprocessing (mostly Sellafield)
134
2.06
Reprocessing and Chernobyl
125
2.7
Reprocessing, mostly Cap de la Hague
99
213 000
Reproducing – large pulse from Sellafield beginning in 1994. Some weapons testing
129
15.7 106
Reprocessing, mostly Cap de la Hague since 1990, some weapons testing
241
14 24 000 6000 88
Weapons testing, reprocessing (mostly Sellafield)
The ‘signature’ reprocessing tracer, used in the earliest studies of Sellafield releases. Has been applied in European coastal waters, the Nordic Seas, Arctic Ocean, and deep North Atlantic In combination with 137Cs, early tracer for Sellafield discharges. The 137Cs/90Sr ratio was used to distinguish reprocessing and weapons sources In combination with 137Cs, this short-lived isotope has been used in the estimation of transit times, e.g., from Sellafield to the northern exit of the Irish Sea Circulation in the English Channel and North Sea, into the Skaggerak and Norwegian coastal waters Surface circulation of European coastal waters, the Nordic Seas, Arctic Ocean, and East and West Greenland Currents Deep water circulation in the Nordic Seas, Arctic Ocean, and Deep Western Boundary Current of the Atlantic Ocean. It has also been measured in European coastal waters and the Gulf of Mexico Not widely used as circulation tracers but often measured in conjunction with other reprocessing radionuclides. Other transuranic elements that have been measured include 241Am and 237Np
Cs
Sr
Cs
Sb
Tc
I
Pu Pu 240 Pu 238 Pu 239
1200
5000
1000
4000
800 3000
600
2000
400 99
1000
Tc
200 0
0 1950
1960
1970
1980
1990
2000
129
I
_ Tc (TBq year 1 ),
_ Cs (TBq year 1)
Cs
137
1400
129
137
99
6000
_ I (TGq year 1 )
Adapted in part from Dahlgaard (1995).
Year Figure 1 Examples of the source functions of reprocessing tracers to the oceans, illustrating some of the differences in magnitude and timing of the releases. 137Cs data, from Gray et al. (1995) and 99Tc data, courtesy of Peter Kershaw, are for liquid discharges from Sellafield. 129I release information is courtesy of G. Raisbeck and F. Yiou, and combines available information for Sellafield from 1966 and Cap de la Hague from 1975. Note the different scales used for the releases of the three isotopes.
Arctic Ocean, reprocessing releases are particularly sensitive tracers of the climatically important deep water formation processes that occur in these regions.
There are several complications associated with this point source tracer introduction, however. Comparison with CFCs indicates some of these difficulties. Within each hemisphere (northern or southern), CFCs are well mixed throughout the troposphere, and the time history of their concentrations is well known. Their entry into the oceans occurs by equilibration with surface waters, and the details of their solubilities as a function of temperature and salinity have been well characterized. The primary complication is that in some areas equilibrium saturation is not reached, and so assumptions must be made about the degree of equilibration. Where they have been necessary, these assumptions appear to be fairly robust. In the case of reprocessing tracers, the discharge amounts have not always been as well known, although this situation has improved continuously. One major difficulty arises in translating a discharge amount, in kg or Bq per year, or per month, to a concentration in sea water some time later. In order to do this, the surface circulation of the coastal regions must be very well known. This circulation is highly variable, on daily, seasonal, interannual and decadal timescales, further
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NUCLEAR FUEL REPROCESSING AND RELATED DISCHARGES
complicating the problem. Also, other sources of the radionuclides under consideration must often be taken into account. Depending on the tracer, these sources may be important and/or numerous, and due to the secrecy involved in many aspects of the operation of nuclear installations the necessary information may not always be available. Other Sources of Anthropogenic Radionuclides
Other sources of radionuclides to the oceans complicate the interpretation of Sellafield and Cap de la Hague tracers to varying extents, depending on the isotope under consideration, the location, and the time. First there is the need to understand the mixing of signals from these two plants. Other sources may include fallout from nuclear weapons tests, uncontrolled releases due to nuclear accidents, dumping on the seabed, other reprocessing plants, atmospheric releases from Sellafield and Cap de la Hague, and unknown sources. Many of these other sources are small compared to the Sellafield and Cap de la Hague releases. The primary complicating source for many isotopes is nuclear weapons test fallout, which peaked in the 1950s and again, more strongly, in 1962–63. This source has been particularly important for 137Cs and 90Sr. The Chernobyl accident in 1986 also released significant amounts of radioactive materials which have themselves found use as oceanographic tracers. In terms of comparison to releases from Sellafield, the Chernobyl accident has been most important with respect to 134Cs and 137Cs. In addition, it has recently been noted that in the case of some nuclides, particularly 137Cs and Pu isotopes, the sediments of the Irish Sea have become a significant ongoing source of tracers to the North Atlantic. Although generally considered a conservative tracer, 137Cs exhibits some affinity for particulate material and a significant amount has accumulated in the sediments around Sellafield. With the continuing reduction of 137Cs activities in liquid effluents from Sellafield, release from the sediments, either by resuspension of the sediments or by reequilibration with the reduced seawater concentrations, has become a relatively large (though still small in an absolute sense compared to the liquid discharges of the 1970s) contributor to the current flux out of the Irish Sea, and will continue to be such for years to come. A further complication of the use of some tracers derived from Sellafield and Cap de la Hague, but one that cannot be lamented, is the continuing reduction of the releases, as well as the radioactive decay of those with the shorter half-lives, such as 137Cs which was released in large quantities over 20 years ago. With respect to many of these complications, and for
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several other reasons, the reprocessing radionuclides attracting the most attention in the oceanographic community today are 99Tc and 129I. Both are long lived and have fairly small relative contributions from weapons testing and other sources. Unlike most radionuclides, their releases increased in the 1990s, and the recent releases of each are largely dominated by a single source: 99Tc by Sellafield and 129I by Cap de la Hague. Significant advances have been and are being made in the measurement of these isotopes, allowing their measurement on smaller sample volumes. The activity of 129I is so low that it is measured by accelerator mass spectrometry. This technique allows measurement of 129I on 1 liter seawater samples. Advances in 99Tc measurement, using both radiochemical and mass-spectrometric (ICP-MS) techniques, are continuing. The primary limitation of 99 Tc studies continues to be the comparative difficulty of its measurement. This is also true to some extent for 129I, for although the sample sizes have been greatly reduced the measurement requires highly specialized technology. A further complication for 129I is its volatility, and the fact that as much as 10% of the reprocessing discharges have been released directly to the atmosphere. Nevertheless, the promise of both tracers is such that these difficulties are likely to be overcome. Regional Setting and Circulation of Reprocessing Discharges
A summary of the regional circulation into which the liquid effluents from Sellafield and Cap de la Hague are released is presented in Figure 2. It is particularly important to note that studies of the reprocessing discharges have contributed greatly to the development of this detailed picture of the regional circulation. Briefly, from the Sellafield site the waste stream is carried north out of the Irish Sea, around the coast of Scotland, through the North Sea, and into the northward-flowing Norwegian Coastal Current (NCC). Transport across the North Sea occurs at various latitudes: some fraction of the reprocessing releases ‘short-circuits’ across the northern part of the North Sea, while some flows farther south along the eastern coast of the UK before turning east and north. Recent studies of the EARP 99Tc pulse from Sellafield have suggested that the rate and preferred transport path of Sellafield releases across the North Sea into the NCC may vary in relation to climatic conditions in the North Atlantic such as the North Atlantic Oscillation (NAO). Radionuclides discharged from the Cap de la Hague reprocessing plant flow north-east through the English Channel and into the North Sea,
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_ 300
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Figure 2 Map of the northern North Atlantic and Nordic Seas, indicating the locations of the Sellafield and Cap de la Hague reprocessing plants (stars), and the major circulation pathways relevant to the discussion of the dispersal of reprocessing wastes. Surface currents are indicated with thin solid lines, and deep waters, (deep overflows from the Nordic Seas, the resulting Deep Western Boundary Current (DWBC) and Labrador Sea Water (LSW)), with heavy dashed lines. Curled arrows indicate the two major areas where deep waters are formed by convective processes. Additional abbreviations: NCC, Norwegian Coastal Current; WSC, West Spitsbergen Current; EGC, East Greenland Current; FBC, Faroe Banks Channel; ISOW, Iceland Scotland Overflow Water; NEADW, Northeast Atlantic Deep Water; DSOW, Denmark Strait Overflow Water.
following the coast and joining the Sellafield releases in the NCC. A small amount of the Sellafield releases and some of the Cap de la Hague releases, which flow closer to the coast, flow east through the Skaggerak and Kattegat to enter the Baltic Sea. The NCC is formed of a mixture of coastal waters (containing the reprocessing tracers) and warm, saline North Atlantic surface waters. Dilution of the reprocessing signal with Atlantic water continues along the northward flow path of the NCC. There is evidence from reprocessing tracers that turbulent eddies between the NCC and Atlantic water result in episodic transport westward into the surface waters of the Norwegian Sea, in addition to a fairly welldefined westward advective transport towards Jan Mayen Island. The NCC branches north of Norway, with one branch, the North Cape (or Nordkap) Current, flowing eastward through the Barents Sea and thence into the Kara Sea and Arctic Ocean, and the remainder flowing north and west as part of the West Spitsbergen Current (WSC). This latter flow branches in the Fram Strait west of Spitsbergen, with some recirculation to the west and south joining the
southward flowing East Greenland Current (EGC), and the remainder entering the Arctic Ocean. The majority of surface outflow from the Arctic is through the Fram Strait into the EGC, thus the bulk of the reprocessing nuclides entering the Arctic Ocean will eventually exit to the Nordic (Greenland, Iceland, and Norwegian) Seas and the North Atlantic. The presence of high concentrations of radionuclides derived from reprocessing in the surface waters of the Barents and Nordic Seas is also an indication of their utility as tracers of deep-water ventilation and formation in these regions, as discussed below.
The Use of Reprocessing Releases in Oceanography Historical Background
The first papers reporting measurements of Sellafield-derived 137Cs in coastal waters appeared in the early 1970s. Much of this early work arose from monitoring efforts by the Division of Fisheries
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NUCLEAR FUEL REPROCESSING AND RELATED DISCHARGES
Research of the UK’s Ministry of Agriculture, Fisheries, and Food, which holds a joint oversight role on the Sellafield discharges and which continues to this day, now as the Centre for Environment, Fisheries and Aquaculture Science, to be a leader in studies of the distribution and oceanographic application of reprocessing radionuclides. In the late 1970s scientists at the Woods Hole Oceanographic Institution published the first of a number of papers detailing the unique utility of the reprocessing tracers and demonstrating their application. Leading studies of reprocessing tracers in the oceans have been undertaken by researchers in numerous other countries affected by the radiological implications of the releases, including France, Germany, Denmark, Canada, Norway, and the former Soviet Union. Coastal and Surface Circulation
A great deal of work has been published using the documented releases of radioisotopes from Sellafield and Cap de la Hague to study the local circulations of the Irish Sea, North Sea, and English Channel. Early studies of the Sellafield releases examined a variety of isotopes, including 134Cs, 137Cs, 90Sr, and Pu isotopes. Most attention focused on 137Cs and 90Sr, and their activity ratio, because: (1) the releases of these two isotopes were well documented, (2) they had been studied extensively since the 1950s and 1960s in weapons test fallout; and (3) the 137Cs/90Sr ratio in reprocessing releases (particularly those from Sellafield) was significantly higher than in global fallout and thus could be used to distinguish the sources of these isotopes in a given water sample. The use of the 134 Cs/137Cs ratio enabled estimates of transit times, assuming the initial ratio in the releases was constant and making use of the short half-life of 134Cs. The short-lived isotope 125Sb has been used as a specific tracer of the circulation of Cap de la Hague discharges through the English Channel and North Sea, into the Baltic and the Norwegian Coastal Current. Summaries of transit times and dilution factors for the transport of Sellafield and Cap de la Hague discharges to points throughout the North Sea, Norwegian Coastal Current, Barents and Kara Seas, Greenland Sea, and East and West Greenland Currents have been published in recent reviews. Numbers are not included in this article because they are currently under revision. In terms of transport to the NCC, where the two waste streams are generally considered to merge, the consensus has been that the transit time from Sellafield to about 601N is three to four years, and that from Cap de la Hague to the same area is one to two years. Compilations of ‘transfer factors,’ which relate observed concentrations to the discharge
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amounts, and factor in the transit times, suggest that the two reprocessing waste streams meet in approximately equal proportions in the NCC. In other words, if Sellafield and Cap de la Hague released equal amounts of a radionuclide, with the Cap de la Hague release two years later, they would make equal contributions to the NCC. Recently, detailed studies have been undertaken of the dispersal of the EARP 99Tc pulse from Sellafield, which began in 1994. These studies have suggested substantially shorter circulation times for Sellafield releases to northern Scottish coastal waters and across the North Sea to the NCC than previously accepted. For instance, the 99Tc pulse reached the NCC within 2.5 years, rather than three to four. It has been suggested that this is a real difference between sampling periods, resulting from climatically induced circulation changes in the North Atlantic. Some of the difference may also be related to the fact that much more detailed data, both in terms of seawater sampling and regarding the releases, are available on this recent event. The continuing passage of the EARP 99Tc signal promises to be very useful in the Nordic Seas and Arctic Ocean as well. Deep-water Formation in the North Atlantic and Arctic Oceans
In addition to surface water flows, deep-water formation processes within the Arctic Ocean and Nordic Seas have been elucidated through the study of reprocessing releases. In a classic presentation of reprocessing tracer data, Livingston showed that the surface water distribution of 137Cs in the Nordic Seas in the early 1980s was marked by high concentrations at the margins of the seas, highlighting the delivery of the isotope in the northward-flowing NCC and WSC to the west and the return flow in the southward-flowing EGC to the east. The opposite distribution was found in the deep waters, with higher concentrations in the center of the Greenland Basin than at the margins, as a result of the ventilation of the Greenland Sea Deep Water by deep convective processes in the center of the gyre. In the Arctic Ocean, elevated 137Cs and 90Sr concentrations and 137Cs/90Sr ratios at 1500 m at the LOREX ice station near the North Pole in 1979 indicated that deep layers of the Arctic Ocean were ventilated from the shelves. Similar observations in deep water north of Fram Strait provided early evidence suggesting a contribution of dense brines from the Barents Sea shelf to the bottom waters of the Nansen Basin. Reprocessing tracers, particularly those, like 129I and 99Tc, which have only a small contribution from other sources such as weapons
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fallout, hold great promise for illuminating eastern Arctic Ocean deep water ventilation processes from the Barents and Kara Sea shelves. The NCC delivers reprocessing tracers to the Barents and Kara relatively rapidly (B 5 years) and more importantly in high concentration. With high tracer concentrations in the area of interest as a source water, the reprocessing tracers may be particularly sensitive tracers of a contribution of dense shelf waters to the deep Arctic Ocean. In addition to the deep waters formed through convection in the polar regions (most notably in the Greenland Sea) which fill the deep basins of the Nordic Seas, intermediate waters are formed which subsequently overflow the sills between Greenland, Iceland, and Scotland and ventilate the deep North Atlantic. The presence of 137Cs and 90Sr from Sellafield was reported in the overflow waters immediately south of the Denmark Straits sampled during the Transient Tracers in the Ocean (TTO) program in 1981. Later, it was demonstrated that reprocessing cesium and strontium could be distinguished in the deep waters as far as TTO Station 214, off the Grand Banks of Newfoundland. No samples were taken for reprocessing radionuclides further south as part of that study, but it was clear in retrospect that the reprocessing signal had traveled even further in the Deep Western Boundary Current (DWBC). Recent work on 129I has highlighted the utility of reprocessing radionuclides as tracers of northern source water masses and the DWBC of the Atlantic. Profiles in stations south of the overflows show much clearer tracer signals for 129I than for the CFCs, and 129 I has been detected in the DWBC as far south as Cape Hatteras. As with the cesium studies of the 1980s, it is likely that sampling further south will reveal the tracer there as well.
Conclusions Releases of radionuclides from the nuclear fuel-reprocessing plants at Sellafield and Cap de la Hague have provided tracers for detailed studies of the circulations of the local environment into which they are released, namely the Irish Sea, English Channel, and North Sea, and for the larger-scale circulation processes of the North Atlantic and Arctic Oceans. These tracers have very different source functions for their introduction into the oceans compared to other widely used anthropogenic tracers, and in some cases compared to each other. The fact that they are released at point sources makes them highly specific tracers of several interesting processes in ocean
circulation, but the nature of the releases has complicated their quantitative interpretation to some extent. Recent advances have been made in the measurement of 129I and 99Tc, long-lived tracers whose releases have increased in recent years and which experience little complication from other sources. These two tracers hold great promise for elucidating deep water formation and ventilation processes in the North Atlantic, Nordic Seas, and Arctic Ocean in the years to come.
See also Arctic Ocean Circulation. CFCs in the Ocean. North Sea Circulation. Radioactive Wastes. Water Types and Water Masses.
Further Reading Aarkrog A, Dahlgaard H, Hallstadius L, Hansen H, and Hohm E (1983) Radiocaesium from Sellafield effluents in Greenland waters. Nature 304: 49--51. Dahlgaard H (1995) Transfer of European coastal pollution to the Arctic: radioactive tracers. Marine Pollution Bulletin 31: 3--7. Gray J, Jones SR, and Smith AD (1995) Discharges to the environment from the Sellafield Site, 1951–1992. Journal of Radiological Protection 15: 99--131. Jefferies DF, Preston A, and Steele AK (1973) Distribution of caesium-137 in British coastal waters. Marine Pollution Bulletin 4: 118--122. Kershaw P and Baxter A (1995) The transfer of reprocessing wastes from north-west Europe to the Arctic. Deep-Sea Research II 42: 1413--1448. Kershaw PJ, McCubbin D, and Leonard KS (1999) Continuing contamination of north Atlantic and Arctic waters by Sellafield radionuclides. Science of Total Environment 237/238: 119--132. Livingston HD (1988) The use of Cs and Sr isotopes as tracers in the Arctic Mediterranean Seas. Philosophical Transactions of the Royal Society of London A 325: 161--176. Livingston HD, Bowen VT, and Kupferman SL (1982) Radionuclides from Windscale discharges I: non-equilibrium tracer experiments in high-latitude oceanography. Journal of Marine Research 40: 253--272. Livingston HD, Bowen VT, and Kupferman SL (1982) Radionuclides from Windscale discharges II: their dispersion in Scottish and Norwegian coastal circulation. Journal of Marine Research 40: 1227--1258. ¨ stlund HG (1985) Livingston HD, Swift JH, and O Artificial radionuclide tracer supply to the Denmark Strait Overflow between 1972 and 1981. Journal of Geophysical Research 90: 6971--6982.
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OCEAN BIOGEOCHEMISTRY AND ECOLOGY, MODELING OF N. Gruber, Institute of Biogeochemistry and Pollutant Dynamics, ETH Zurich, Switzerland S. C. Doney, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Modeling has emerged in the last few decades as a central approach for the study of biogeochemical and ecological processes in the sea. While this development was facilitated by the fast development of computer power, the main driver is the need to analyze and synthesize the rapidly expanding observations, to formulate and test hypotheses, and to make predictions how ocean ecology and biogeochemistry respond to perturbations. The final aim, prediction, has gained in importance recently as scientists are increasingly asked by society to investigate and assess the impact of past, current, and future human actions on ocean ecology and biogeochemistry. The impact of the carbon dioxide (CO2) that humankind has emitted and will continue to emit into the atmosphere for the foreseeable future is currently, perhaps, the dominant question facing the marine biogeochemical/ecological research community. This impact is multifaceted, and includes both direct (such as ocean acidification) and indirect effects that are associated with the CO2-induced climate change. Of particular concern is the possibility that global climate change will lead to a reduced capacity of the ocean to absorb CO2 from the atmosphere, so that a larger fraction of the CO2 emitted into the atmosphere remains there, further enhancing global warming. In such a positive feedback case, the expected climate change for a given CO2 emission will be larger relative to a case without feedbacks. Marine biogeochemical/ecological models have played a crucial role in elucidating and evaluating these processes, and they are increasingly used for making quantitative predictions with direct implications for climate policy. There are many other marine biogeochemical and/ or ecological problems related to human activities, for which models play a crucial role assessing their importance and magnitude and devising possible
solutions. These include, for example, coastal eutrophication, overfishing, and dispersion of invasive species. The use of marine biogeochemical/ecological models is now so pervasive that practically every field of oceanography is on this list. The aim of this article is to provide an introduction and overview of marine biogeochemical and ecological modeling. Given the breadth of modeling approaches in use today, this overview can by design not be inclusive and authoritative. We rather focus on some basic concepts and provide a few illustrative applications. We start with a broader description of the marine biogeochemical/ecological challenge at hand, and then introduce basic concepts used for biogeochemical/ecological modeling. In the final section, we use a number of examples to illustrate some of the core modeling approaches.
The Marine Ecology and Biogeochemistry Challenge The complexity of the ocean biogeochemical/ecological problem is daunting, as it involves a complex interplay among biology, physical variability of the oceanic environment, and the interconnected cycles of a large number of bioactive elements, particularly those of carbon, nitrogen, phosphorus, oxygen, silicon, and iron (Figure 1). Furthermore, the ocean is an open system that exchanges mass and many elements with the surrounding realms, such as the atmosphere, the land, and the sediments. The engine that sets nearly all of these cycles into motion is the photosynthetic fixation of dissolved inorganic carbon and many other nutrient elements into organic matter by phytoplankton in the illuminated upper layers of the ocean (euphotic zone). The net rate of this process (i.e., net primary production) is distributed heterogeneously in the ocean, primarily as a result of the combined limitation of nutrients and light. This results in similar heterogeneity in surface chlorophyll, a direct indicator of the amount of phytoplankton biomass (Figure 2(a)). The large-scale surface nutrient distributions, in turn, reflect the balance between biological removal and the physical processes of upwelling and mixing that transport subsurface nutrient pools upward into the euphotic zone (Figure 2(b)). In addition to the traditional macronutrients (nitrate, phosphate, silicate etc.), growing evidence shows that iron limitation is
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Atmosphere N2
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Figure 1 Schematic diagram of a number of key biogeochemical cycles in the ocean and their coupling. Shown are the cycles of carbon, oxygen, phosphorus, nitrogen, and silicon. The main engine of most biogeochemical cycles in the ocean is the biological production of organic matter in the illuminated upper ocean (euphotic zone), part of which sinks down into the ocean’s interior and is then degraded back to inorganic constituents. This organic matter cycle involves not only carbon, but also nitrogen, phosphorus, and oxygen, causing a tight linkage between the cycles of these four elements. Since some phytoplankton, such as diatoms and coccolithophorids, produce shells made out of amorphous silicon and solid calcium carbonate, the silicon and CaCO3 cycles also tend to be closely associated with the organic matter cycle. These ocean interior cycles are also connected to the atmosphere through the exchange of a couple of important gases, such as CO2, oxygen, and nitrous oxide (N2O). In fact, on timescales longer than a few decades, the ocean is the main controlling agent for the atmospheric CO2 content.
a key factor governing net primary production in many parts of the ocean away from the main external sources for iron, such as atmospheric dust deposition and continental margin sediments. The photosynthesized organic matter is the basis for a complex food web that involves both a transfer of the organic matter toward higher trophic levels as well as a microbial loop that is responsible for most of the breakdown of this organic matter back to its inorganic constituents. A fraction of the synthesized organic matter (about 10–20%) escapes degradation in the euphotic zone and sinks down into the dark aphotic zone, where it fuels the growth of microbes and zooplankton that eventually also remineralize most of this organic matter back to its inorganic constituents. The nutrient elements and the dissolved inorganic carbon are then eventually returned back to the surface ocean by ocean circulation and mixing, closing this ‘great biogeochemical loop’. This loop is slightly leaky in that a
small fraction of the sinking organic matter fails to get remineralized in the water column and is deposited onto the sediments. Very little organic matter escapes remineralization in the sediments though, so that only a tiny fraction of the organic matter produced in the surface is permanently removed from the ocean by sediment burial. In the long-term steady state, this loss of carbon and other nutrient elements from the ocean is replaced by the input from land by rivers and through the atmosphere. Several marine phytoplankton and zooplankton groups produce hard shells consisting of either mineral calcium carbonate (CaCO3) or amorphous silica, referred to as ‘opal’ (SiO2). The most important producers of CaCO3 are coccolithophorids, while most marine opal stems from diatoms. Both groups are photosynthetic phytoplankton, highlighting the extraordinary importance of this trophic group for marine ecology and biogeochemistry. Upon the death of the mineral-forming organisms, these minerals
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while respiration and bacterial degradation releases CO2 and consumes O2. As a result, one tends to find an inverse relationship in the oceanic distribution of these two gases. CO2 also sets itself apart from O2 in that it reacts readily with seawater. In fact, due to the high content of alkaline substances in the ocean, this reaction is nearly complete, so that only around 1% of the total dissolved inorganic carbon in the ocean exists in the form of CO2. The exchange of CO2 across the air–sea interface is a prime question facing ocean biogeochemical/ ecological research, particularly with regard to the magnitude of the oceanic sink for anthropogenic CO2. The anthropogenic CO2 sink occurs on top of the natural air–sea CO2 fluxes characterized by oceanic uptake in mid-latitudes and some high latitudes, and outgassing in the low latitudes and the Southern Ocean. This distribution of the natural CO2 flux is the result of an interaction between the exchange of heat between the ocean and the atmosphere, which affects the solubility of CO2, and the great biogeochemical loop, which causes an uptake of CO2 from the atmosphere in regions where the downward flux of organic carbon exceeds the upward supply of dissolved inorganic carbon, and an outgassing where the balance is the opposite. This flux has changed considerably over the last two centuries in response to the anthropogenically driven increase in atmospheric CO2 that pushes additional CO2 from the atmosphere into the ocean. Nevertheless, the pattern of the resulting contemporary air–sea CO2 flux (Figure 2(c)) primarily still reflects the flux pattern of the natural CO2 fluxes, albeit with a global integral flux into the ocean reflecting the oceanic uptake of anthropogenic CO2. Given the central role of the great biogeochemical loop, any modeling of marine biogeochemical/ecological processes invariably revolves around the modeling of all the processes that make up this loop. As this loop starts with the photosynthetic production of organic matter by phytoplankton, biogeochemical modeling is always tightly interwoven with the modeling of marine ecology, especially that of the lower trophic levels. Core questions that challenge marine biogeochemistry and ecology are as follows: 1. What controls the mean concentration and threedimensional (3-D) distribution of bioreactive elements in the ocean? 2. What controls the air–sea balance of climatically important gases, that is, CO2, N2O, and O2? 3. What controls ocean productivity, the downward export of organic matter, and the transfer of organic matter to higher trophic levels?
4. How do ocean biogeochemistry and ecology change in time in response to climate dynamics and human perturbations? Modeling represents a powerful approach to studying and addressing these core questions. Modeling is by no means the sole approach. In fact, integrated approaches that combine observational, experimental, and modeling approaches are often necessary to tackle this set of complex problems.
What Is a Biogeochemical/Ecological Model? At its most fundamental level, a model is an abstract description of how some aspect of nature functions, most often consisting of a set of mathematical expressions. In the biogeochemical/ecological modeling context, a model usually consists of a number of partial differential equations, which describe the time and space evolution of a (limited) number of ecological/biogeochemical state variables. As few of these equations can be solved analytically, they are often solved numerically using a computer, which requires the discretization of these equations, that is, they are converted into difference equations on a predefined spatial and temporal grid.
The Art of Biogeochemical/Ecological Modeling A model can never fully represent reality. Rather, it aims to represent an aspect of reality in the context of a particular problem. The art of modeling is to find the right level of abstraction, while keeping enough complexity to resolve the problem at hand. That is, marine modelers often follow the strategy of Occam’s razor, which states that given two competing explanations, the one that is simpler and makes fewer assumptions is the more likely to be correct. Therefore, a typical model can be used only for a limited set of applications, and great care must be used when a model is applied to a problem for which it was not designed. This is especially true for biogeochemical/ecological models, since their underlying mathematical descriptions are for the most part not based on first principles, but often derived from empirical relationships. In fact, marine biogeochemical/ecological modeling is at present a data-limited activity because we lack data to formulate and parametrize key processes and/or to evaluate the model predictions. Further, significant simplifications are often made to make the problem more tractable. For example, rather than treating
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OCEAN BIOGEOCHEMISTRY AND ECOLOGY, MODELING OF
individual organisms or even species, model variables often aggregate entire functional groups into single boxes (e.g., photosynthetic organisms, grazers, and detritus decomposers), which are then simulated as bulk concentrations (e.g., mol C m 3 of phytoplankton). Despite these limitations, models allow us to ask questions about the ocean inaccessible from data or experiments alone. In particular, models help researchers quantify the interactions among multiple processes, synthesize diverse observations, test hypotheses, extrapolate across time – space scales, and predict past and future behavior. A well-posed model encapsulates our understanding of the ocean in a mathematically consistent form.
Biogeochemical/Ecological Modeling Equations and Approaches In contrast with their terrestrial counterparts, models of marine biogeochemical/ecological processes must be coupled to a physical circulation model of some sort to take into consideration that nearly all relevant biological and biogeochemical processes occur either in the dissolved or suspended phase, and thus are subject to mixing and transport by ocean currents. Thus, a typical coupled physical–biogeochemical/ ecological model consists of a set of time-dependent advection, diffusion, and reaction equations: @C þ AdvðCÞ þ DiffðCÞ ¼ SMSðCÞ @t
½1
where C is the state variable to be modeled, such as the concentration of phytoplankton, nutrients, or dissolved inorganic carbon, often in units of mass per unit volume (e.g., mol m 3). Adv(C) and Diff(C) are the contributions to the temporal change in C by advection and eddy-diffusion (mixing), respectively, derived from the physical model component. The term SMS(C) refers to the ‘sources minus sinks’ of C driven by ecological/biogeochemical processes. The SMS term often involves complex interactions among a number of state variables and is provided by the biogeochemical/ecological model component. Marine biogeochemical/ecological models are diverse, covering a wide range of complexities and applications from simple box models to globally 4-D(space and time) coupled physical–biogeochemical simulations, and from strict research tools to climate change projections with direct societal implications. Model development and usage are strongly shaped by the motivating scientific or policy problems as well as the dynamics and time–space scales considered. The
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complexity of marine biogeochemical/ecological models can be organized along two major axes (Figure 3): the physical complexity, which determines how the left-hand side of [1] is computed, and the biogeochemical/ecological complexity, which determines how the right-hand side of [1] is evaluated. Due to computational and analytical limitations, there is often a trade-off between the physical and biogeochemical/ecological complexity, so that models of the highest physical complexity are often using relatively simple biogeochemical/ecological models and vice versa (Figure 4). Additional constraints arise from the temporal domain of the integrations. Applications of coupled physical–biogeochemical/ ecological models to paleoceanographic questions require integrations of several thousand years. This can only be achieved by reducing both the physical and the biogeochemical/ecological complexity (Figure 4). At the same time, the continuously increasing computational power has permitted researchers to push forward along both complexity axes. Nevertheless, the fundamental tradeoff between physical and biogeochemical/ecological complexity remains. In addition to the physical and biogeochemical/ ecological complexity, models can also be categorized with regard to their interaction with observations. In the case of ‘forward models’, a set of equations in the form of [1] is integrated forward in time given initial and boundary conditions. A typical forward problem is the prediction of the future state of ocean biogeochemistry and ecology for a certain evolution of the Earth’s climate. The solutions of such forward models are the time–space distribution of the state variables as well as the implied fluxes. The expression ‘inverse models’ refers to a broad palette of modeling approaches, but all of them share the goal of optimally combining observations with knowledge about the workings of a system as embodied in the model. Solutions to such inverse models can be improved estimates of the current state of the system (state estimation), improved estimates of the initial or boundary conditions, or an optimal set of parameters. A typical example of an inverse model is the optimal determination of ecological parameters, such as growth and grazing rates, given, for example, the observed distribution of phytoplankton, zooplankton, and nutrients.
Examples Given the large diversity of marine biogeochemical/ ecological models and approaches, no review can do
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Higher complexity
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Sinking particles with aphotic ecosystem determining remineralization
Variable stoichiometry
Sediments passive
Sediments active
Physical complexity
Spatial resolution
System coupling
Physiological modeling
Higher complexity
box
ocean only
1-D
2-D
3-D
ocean/ sea-ice
3-D (eddy resolving)
ocean/ sea-ice/ atmosphere/ land surface
Figure 3 Schematic diagram summarizing the development of complexity in coupled physical–biogeochemical/ecological models. The two major axes of complexity are biogeochemistry/ecology and physics, but each major axis consists of many subaxes that describe the complexity of various subcomponents. Currently existing models fill a large portion of the multidimensional space opened by these axes. In addition, the evolution of models is not always necessarily straight along any given axis, but depends on the nature of the particular problem investigated. N, nutrient; P, phytoplankton; Z, zooplankton; B, bacteria; D, detritus; F, fish.
full justice. We restrict our article here to the discussion of four examples, which span the range of complexities as well as have been important milestones in the evolution of biogeochemical/ecological modeling. Box Models or What Controls Atmospheric Carbon Dioxide?
Our aim here is to develop a model that explains how the great biogeochemical loop controls atmospheric CO2. The key to answering this question is a quantitative prediction of the surface ocean concentration of
CO2, as it is this surface concentration that controls the atmosphere–ocean balance of CO2. The simplest models used for such a purpose are box models, where the spatial dimension is reduced to a very limited number of discrete boxes. Such box models have played an important role in ocean biogeochemical/ ecological modeling, mostly because their solutions can be readily explored and understood. However, due to the dramatic reduction of complexity, there are also important limitations, whose consequences one must keep in mind when interpreting the results. Box models can be formally derived from the tracer conservation eq [1] by integrating over
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OCEAN BIOGEOCHEMISTRY AND ECOLOGY, MODELING OF
95
n
Higher complexity
ai
m
o ld
a
or
e mp tim d te U an CP int g o in p as id re r gr c In pe
Biogeochemical/ecological complexity
Ecological case studies (e.g., test-beds) short-term integrations
Climate-change studies century-timescale integrations
Temporal evolution
Physical−biological coupling studies short-term integrations
Geochemical applications long-term integrations
Physical complexity
Higher complexity
Figure 4 Schematic illustrating the relationship between the physical (abscissa) and the biogeochemical/ecological complexity (ordinate) of different typical applications of coupled physical–biogeochemical/ecological models. Given computational and analytical constraints, there is often a trade-off between the two main complexities. The thin lines in the background indicate the CPU time required for the computation of a problem over a given time period.
the volume of the box, V, and by applying Gauss’ (divergence) theorem, which states that the volume integral of a flux is equal to the flux in normal direction across the boundary surfaces, S. The resulting integral form of [1] is: Z
I
@C dV þ ðAdvn ðCÞ þ Diff n ðCÞÞ dS @t Z ¼ SMSðCÞ dV
½2
where Advn(C) and Diffn(C) are the advective and diffusive transports in the normal direction across S. If we assume that the concentration of tracer C as well as the SMS term within the box are uniform, [2]
can be rewritten as V
dC X þ Ti;out C Ti;in Ci þ ni ðC Ci Þ dt i ¼ V SMSðCÞ
½3
where C is the mean concentration within the box. We represented the advective contributions as products of a mass transport, T (dimensions volume time 1), with the respective upstream concentration, separately considering the mass transport into the box and out of the box across each interface i, that is, Ti,in and Ti,out. The eddy-diffusion contributions are parametrized as products of a mixing coefficient, ni (dimension volume time 1), with the gradient
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across each interface ðC Ci Þ. For simplicity, we subsequently drop the overbar, that is, use C instead of C. The simplest box-model representation of the great biogeochemical loop is a two-box model, wherein the ocean is divided into a surface and a deep box, representing the euphotic and aphotic zones, respectively (Figure 5(a)). Mixing and transport between the two boxes is modeled with a mixing term, n. The entire complexity of marine ecology and biogeochemistry in the euphotic zone is reduced to a single term, F, which represents the flux of organic matter out of the surface box and into the deep box, where the organic matter is degraded back to inorganic constituents. Let us consider the deep-box balance for a dissolved inorganic bioreactive element, Cd (such as nitrate, phosphate, or dissolved inorganic carbon): Vd
dCd þ n ðCd Cs Þ ¼ F dt
½4
where Vd is the volume of the deep box and Cs is the dissolved inorganic concentration of the surface box. The steady-state solution of [4], that is, the solution when the time derivative of Cd vanishes (dCd/dt ¼ 0), n ðCd Cs Þ ¼ F
½5
represents the most fundamental balance of the great biogeochemical loop. It states that the net upward transport of an inorganic bioreactive element by physical mixing is balanced by the downward transport of this element in organic matter. This steady-state balance [5] has several important applications. For example, it permits us to estimate the downward export of organic matter by simply
analyzing the vertical gradient and by estimating the vertical exchange. The balance also states that the magnitude of the vertical gradient in bioreactive elements is proportional to the strength of the downward flux of organic matter and inversely proportional to the mixing coefficient. Since nearly all oceanic inorganic carbon and nutrients reside in the deep box, that is, Vd Cd cVs Cs , one can assume, to first order, that Cd is largely invariant, so that the transformed balance for the surface ocean concentration Cs ¼ Cd
F n
½6
reveals that the surface concentration depends primarily on the relative magnitude of the organic matter export to the magnitude of vertical mixing. Hence, [6] provides us with a first answer to our challenge: it states that, for a given amount of ocean mixing, the surface ocean concentration of inorganic carbon, and hence atmospheric CO2, will decrease with increasing marine export production. Relationship [6] also states that for a given magnitude of marine export production, atmospheric CO2 will increase with increased mixing. While very powerful, the two-box model has severe limitations. The most important one is that this model does not consider the fact that the deep ocean exchanges readily only with the high latitudes, which represent only a very small part of the surface ocean. The exchange of the deep ocean with the low latitudes is severly limited because diapcyncal mixing in the ocean is small. Another limitation is that the onebox model does not take into account that the nutrient concentrations in the high latitudes tend to be much higher than in the low latitudes (Figure 2(c)).
(a)
(b) High-lat
Surface Vs
Low-lat
Vh
Cs
Cl Vl
Ch T
ν
Cd
Vd
Deep
Three-box model
Two-box model
Φ
Φl
hd Φh Cd Vd
Deep
Figure 5 Schematic representation of the great biogeochemical loop in box models: (a) two-box model and (b) three-box model. C, concentrations; V, volume; n, exchange (mixing) coefficient; F, organic matter export fluxes; T, advective transport.
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OCEAN BIOGEOCHEMISTRY AND ECOLOGY, MODELING OF
= +145
(a)
97
(c) = −120
(b)
Atmospheric pCO2 (ppm)
280
160
425
0
33
0
26
33
0
0
33
33
Nitrate balance (mmol−3) Window open
Preindustrial
Window closed
Figure 6 Schematic illustration of the impact of the high-latitude (Southern Ocean) window on atmospheric CO2. In (a) the highlatitude window is open, high-latitude nitrate is high, and atmospheric CO2 attains 425 ppm. Panel (b) represents the preindustrial ocean, where the window is half open, nitrate in the high latitudes is at intermediate levels, and atmospheric CO2 is 280 ppm. In (c), the high-latitude window is closed, high-latitude nitrate is low, and atmospheric CO2 decreases to about 160 ppm.
An elegant solution is the separation of the surface box into a high-latitude box, h, and a low-latitude box, l (Figure 5(b)). Intense mixing is assumed to occur only between the deep and the high-latitude boxes, nhd . A large-scale transport, T, is added to mimic the ocean’s global-scale overturning circulation. As before, the complex SMS terms are summarized by Fl and Fh. In steady state, the surface concentrations are given by Fl T Fh þ Fl Ch ¼ Cd T þ nhd Cl ¼ Cd
½7
which are structurally analogous to [6], but emphasize the different dynamics setting balances in the low and high latitudes. With nhd cT and FlEFh, [7] explains immediately why nutrient concentrations tend to be higher in high latitudes than in low latitudes (Figure 2(b)). In fact, writing [7] out for both nitrate ðNO 3 Þ and dissolved inorganic carbon (DIC), and using the observation that surface NO 3 in the low latitudes is essentially zero (Figure 2(c)), the DIC equations are DICl ¼ DICd rC:N ½NO 3 d DICh ¼ DICd
Fh þ rC:N T ½NO 3 d T þ nhd
½8
where rC:N is the carbon-to-nitrogen ratio of organic matter. Assuming that the deep-ocean nitrate concentration is time invariant, the analysis of [8] reveals that the low-latitude DIC concentration, DICl,
is more or less fixed, while the high-latitude DIC concentration, DICh, can be readily altered by changes in either the high-latitude export flux, Fh, or high-latitude mixing, nhd . Furthermore, if one considers the fact that the rapid communication between the high-latitude ocean and the deep ocean makes the high latitudes the primary window into the oceanic reservoir of inorganic carbon, it becomes clear that it must be the high latitudes that control atmospheric CO2 on the millenial timescales where the steadystate approximation is justified. This high-latitude dominance in controlling atmospheric CO2 is illustrated in Figure 6, which demonstrates that atmospheric CO2 can vary between about 160 and 425 ppm by simply opening or closing this highlatitude window. The exact magnitude of the atmospheric CO2 change depends, to a substantial degree, on the details of the model, but the dominance of high-latitude processes in controlling atmospheric CO2 is also found in much more complex and spatially explicit models. As a result, a change in the carbon cycle in the high latitudes continues to be the leading explanation for the substantially lower atmospheric CO2 concentrations during the ice ages. NP and NPZ Models, or What Simple Ecosystem Models Can Say about Oceanic Productivity
So far, we have represented the ecological, chemical, and physical processes that control the production of organic matter and its subsequent export with a single parameter, F. Clearly, in order to assess how marine biology responds to climate change and other perturbations, it is necessary to resolve these
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processes in much more detail. Two fundamentally different approaches have been developed to model such ecological processes: concentration-based models and individually based models (IBMs). In the latter, the model’s equations represent the growth and losses of individual organisms, and the model then simulates the evolution of a large number of these individuals through space and time. This approach is most commonly used for the modeling of organisms at higher trophic levels, where life cycles play an important role (e.g., models of zooplankton, fish, and marine mammals). In contrast, concentration-based models are almost exclusively used for the modeling of the lower trophic levels, in particular to represent the interaction of light, nutrient, and grazing in controlling the growth of phytoplankton, that is, marine primary production. The simplest such model considers just one limiting nutrient, N, and one single phytoplankton group, P (see also Figure 7(a)): N P lP P KN þ N N P þ mP l P P SMSðNÞ ¼ Vmax KN þ N SMSðPÞ ¼ Vmax
½9
where the first term on the right-hand side of the phytoplankton eqn [9] is net phytoplankton growth (equal to net primary production) modeled here as a function of a nutrient-saturated growth rate Vmax, and a hyperbolic dependence on the in situ nutrient concentration (Monod-type), with a single parameter, the half-saturation constant, KN. The second term is the net loss due to senescence, viral infection, and grazing, modeled here as a linear process with a loss rate lP . A fraction mP of the phytoplankton loss is assumed to be regenerated inside the euphotic zone, while a fraction (1 mP) is exported to depth (Figure 7(a)). These two parametrizations for phytoplankton growth and losses reflect a typical situation for marine ecosystem models in that the growth terms are substantially more elaborate, reflecting the interacting influence of various controlling parameters, whereas the loss processes are highly simplified. This situation also reflects the fact that the processes controlling the growth of marine organisms are often more amenable to experimental studies, while the loss processes are much harder to investigate with careful experiments. When the SMS terms of [10] are inserted into the full tracer conservation eqn [1], the resulting equations form a set of coupled partial differential equations with a number of interesting consequences in steady state as shown in Figure 8(a). Below a certain nutrient concentration threshold, phytoplankton
growth is smaller than its loss term, so that the resulting steady-state solution is P ¼ 0, that is, no phytoplankton. Above this threshold, the abundance of phytoplankton (and hence primary production) increases with increasing nutrient supply in a linear manner. Phytoplankton is successful in reducing the dissolved inorganic nutrient concentration to low levels, so that the steady state is characterized by most nutrients residing in organic form in the phytoplankton pool. This NP model reflects a situation where marine productivity is limited by bottom-up processes, that is, the supply of the essential nutrients. Comparison with the nutrient and phytoplankton distribution shown in (Figure 2(b)) shows that this model could explain the observations in the subtropical open ocean, where nutrients are indeed drawn down to very low levels. However, the NP model would also predict relatively high P levels to go along with the low nutrient levels, which is clearly not observed (Figure 2(a)). In addition, this NP model also fails clearly to explain the high-nutrient regions of the high latitudes. However, we have so far neglected the limitation by light, as well as the impact of zooplankton grazing. The addition of a zooplankton compartment to the NP model (Figure 7(b)) dramatically shifts the nutrient allocation behavior (Figure 8(b)). In this case, as the nutrient loading increases, the phytoplankton abundance gets capped at a certain level by zooplankton grazing. Due to the reduced levels of biomass, the phytoplankton is then no longer able to consume all nutrients at high-nutrient loads, so that an increasing fraction of the supplied nutrients remains unused. This top-down limitation situation could therefore, in part, explain why macronutrients in certain regions remain untapped. Most recent research suggests that micronutrient (iron) limitation, in conjunction with zooplankton grazing, plays a more important role in causing these highnutrient/low chlorophyll regions. Another key limitation is light, which has important consequences for the seasonal and depth evolution of phytoplankton. Global 3-D Modeling of Ocean Biogeochemistry/ Ecology
The coupling of relatively simple NPZ-type models to global coarse-resolution 3-D circulation models has proven to be a challenging task. Perhaps the most important limitation of such NPZ models is the fact that all phytoplankton in the ocean are represented by a single phytoplankton group, which is grazed upon by a single zooplankton group. This means that the tiny phytoplankton that dominate the relatively nutrient-poor central gyres of the ocean, and the
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99
(a) N−P model
Mortality and respiration
Phytoplankton (P )
P · P · P
Export
Remineralization
Phytoplankton growth
Nutrient supply
Γ(N )
N KN + N
P · Vmax ·
DIN + DON (N )
P · P
(1 – P ) · P · P
(b) N−P−Z model
Phytoplankton growth
Phytoplankton (P)
P · P · P P · P Export
Remineralization
Γ (N )
Nutrient supply
Remineralization
DIN + DON (N)
N KN + N
Mortality and respiration
Grazing
P · Vmax ·
g·Z·P KP
(1 – P ) · P · P
Zooplankton (Z )
(1 – Z ) ·
g·Z·P + Z · Z KP
Export
Z ·
(1 – Z ) · Mortality and respiration and egestion (1 – Z ) ·
(1 – Z ) ·
g·Z·P + Z ·Z KP
g·Z·P + Z · Z KP
Figure 7 A schematic illustration of (a) a nitrate–phytoplankton model and (b) a nitrate–phytoplankton–zooplankton model. The mathematical expressions associated with each arrow represent typical representations for how the individual processes are parametrized in such models. Vmax, maximum phytoplankton growth rate; KN, nutrient half saturation concentration; g , maximum zooplankton growth rate; KP, half saturation concentration for phytoplankton grazing; gZ, zooplankton assimilation efficiency; lP, phytoplankton mortality rate; lZ, zooplankton mortality rate; mP, fraction of dead phytoplankton nitrogen that is remineralized in the euphotic zone; mZ, as mP, but for zooplankton. Adapted from Sarmiento JL and Gruber N (2006) Ocean Biogeochemical Dynamics, 526pp. Princeton, NJ: Princeton University Press.
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(a) 0.03
(b) 10
NP-model
NPZ-model
0.02 NT
=
Nn
P, Nn + P (mmol m–3)
P, Nn + P (mmol m–3)
8 +
P
0.01 Nn
P
6
Nn
4
2 Z
0
0
0.01
0.02 NT (mmol m–3)
0.03
0
P 0
2
4 6 NT (mmol m–3)
8
10
Figure 8 Nitrogen in phytoplankton (P), zooplankton (Z), and nitrate (Nn) as a function of total nitrogen. (a) Results from a twocomponent model that just includes N and P. (b) Results from a three-component model that is akin to that shown in Figure 7(b). The vertical axis is concentration in each of the components plotted as the cumulative amount. Adapted from Sarmiento JL and Gruber N (2006) Ocean Biogeochemical Dynamics, 526pp. Princeton, NJ: Princeton University Press.
large phytoplankton that dominate the highly productive upwellling regions are represented with a single group, whose growth characteristics cannot simultaneously represent both. An additional challenge of the large-scale coupling problem is the interaction of oceanic circulation/mixing with the marine ecological and biogeochemical processes. Often, small deficiencies in the physical model get amplified by the ecological/biogeochemical model, so that the resulting fields of phytoplankton abundance may diverge strongly from observations. The latter problem is addressed by improving critical aspects of the physical model, such as upper ocean mixing, atmospheric forcing, resolution, and numerical tracer transport algorithms, while the first problem is addressed by the consideration of distinct plankton functional groups. These functional groups distinguish themselves by their different size, nutrient requirement, biogeochemical role, and several other characteristics, but not necessarily by their taxa. Currently existing models typically consider the following phytoplankton functional groups: small phytoplankton (nano- and picoplankton), silicifying phytoplankton (diatoms), calcifying phytoplankton (coccolithophorids), N2fixing phytoplankton, large (nonsilicifying) phytoplankton (e.g., dinoflagellates and phaeocystis), with the choice of which group to include being driven by the particular question at hand. The partial differential equations for the different phytoplankton functional groups are essentially the same and follow a structure similar to [9], but are differentiated by varying growth parameters, nutrient/light requirements, and susceptibility to grazing. At present, most applications include between five and 10
phytoplankton functional groups. Some recent studies have been exploring the use of several dozen functional groups, whose parameters are generated stochastically with some simple set of rules that are then winnowed via competition for resources among the functional groups. Since the different phytoplankton functional groups are also grazed differentially, different types of zooplankton generally need to be considered as well, though fully developed models with diverse zooplankton functional groups are just emerging. An alternative is to use a single zooplankton group, but with changing grazing/ growth characteristics depending on its main food source. A main advantage of models that build on the concept of plankton functional groups is their ability to switch between different phytoplankton community structures and to simulate their differential impact on marine biogeochemical processes. For example, the nutrient poor subtropical gyres tend to be dominated by nano- and picoplankton, whose biomass tends to be tightly capped by grazing by very small zooplankton. This prevents the group forming blooms, even under nutrient-rich conditions. By contrast, larger phytoplankton, such as diatoms, can often escape grazing control (at least temporarily) and form extensive blooms. Such phytoplankton community shifts are essential for controlling the export of organic matter toward the abyss, since the export ratio, that is, the fraction of net primary production that is exported, tends to increase substantially in blooms. Results from the coupling of such a multiple functional phytoplankton ecosystem model and of a relatively simple biogeochemical model to a global 3-D circulation model show substantial success in
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(a)
(b)
34.6° N
Nitrate (mmol−3)
Chlorophyll (mg m−3)
34.2° N 33.8° N 33.4° N 33.0° N 32.6° N 121° W
0
120° W
0.4 0.8 1.2 1.6
119° W
2
118° W
2.4 2.8 3.2 3.6
121° W
4
15
0.1
120° W
0.17
0.29
119° W
0.5
0.87
118° W
1.5
2.5
4.4
Figure 10 Snapshots of (a) surface nitrate and (b) surface chlorophyll as simulated by a 1-km, that is, eddy-resolving, regional model for the Southern California Bight. The snapshots represent typical conditions in response to a spring-time upwelling event in this region. The model consists of a multiple phytoplankton functional group model that has been coupled to the Regional Ocean Modeling System (ROMS). Results provided by H. Frenzel, UCLA.
simulations). This is currently feasible for only very short periods, so that limited domain models are often used instead. Eddy-resolving Regional Models
Limited domain models, as shown in Figure 10 for the Southern California Bight, that is, the southernmost region of the West Coast of the US, can be run at much higher resolutions over extended periods, permitting the investigation of the processes occurring at the kilometer scale. The most important such processes are meso- and submesoscale eddies and other manifestations of turbulence, which are ubiquitous features of the ocean. The coupling of the same ecosystem/biogeochemical model used for the global model shown in Figure 9 to a 1-km resolution model of the Southern California Bight reveals the spatial richness and sharp gradients that are commonly observed (Figure 10) and that emerge from the intense interactions of the physical and biogeochemical/ecological processes at the eddy scale. Since such mesoscale processes are unlikely to be resolved at the global scale for a while, a current challenge is to find parametrizations that incorporate the net impact of these processes without actually resolving them explicitly.
Future Directions The modeling of marine biogeochemical/ecological processes is a new and rapidly evolving field, so that
predictions of future developments are uncertain. Nevertheless, it is reasonable to expect that models will continue to evolve along the major axes of complexity, including, for example, the consideration of higher trophic levels (Figure 4). One also envisions that marine biogeochemical/ecological models will be increasingly coupled to other systems. A good example are Earth System Models that attempt to represent the entire climate system of the Earth including the global carbon cycle. In those models, the marine biogeochemical/ecological models are just a small part, but interact with all other components, possibly leading to complex behavior including feedbacks and limit cycles. Another major anticipated development is the increasing use of inverse modeling approaches to optimally make use of the increasing flow of observations.
Glossary advection The transport of a quantity in a vector field, such as ocean circulation. anthropogenic carbon The additional carbon that has been released to the environment over the last several centuries by human activities including fossil-fuel combustion, agriculture, forestry, and biomass burning. biogeochemical loop, great A set of key processes that control the distribution of bioreactive elements in the ocean. The loop starts with the photosynthetic production of organic matter in the light-illuminated upper ocean, a fraction of which
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OCEAN BIOGEOCHEMISTRY AND ECOLOGY, MODELING OF
is exported to depth mostly by sinking of particles. During their sinking in the dark deeper layers of the ocean, this organic matter is consumed by bacteria and zooplankton that transform it back to its inorganic constituents while consuming oxidized components in the water (mostly oxygen). The great biogeochemical loop is closed by the physical transport and mixing of these inorganic bioreactive elements back to the near-surface ocean, where the process starts again. This great biogeochemical loop also impacts the distribution of many other elements, particularly those that are particle reactive, such as thorium. coccolithophorids A group of phytoplankton belonging to the haptophytes. They are distinguished by their formation of small mineral calcium carbonate plates called coccoliths, which contribute to the majority of marine calcium carbonate production. An example of a globally significant coccolithophore is Emiliania huxleyi. diatoms A group of phytoplankton belonging to the class Bacillariophyceae. They are one of the most common types of phytoplankton and distinguish themselves through their production of a cell wall made of amorphous silica, called opal. These walls show a wide diversity in form, but usually consist of two symmetrical sides with a split between them. diffusion The transport of a quantity by Brownian motion (molecular diffusion) or turbulence (eddy diffusion). The latter is actually some form of advection, but since its net effect is akin to diffusion, it is often represented as a diffusive process. export production The part of the organic matter formed in the surface layer by photosynthesis that is transported out of the surface layer and into the interior of the ocean by particle sinking, mixing, and circulation, or active transport by organisms. models, concentration based Models, in which the biological state variables are given as numbers per unit volume, that is, concentration. This approach is often used when physical dispersion plays an important role in determining the biological and biogeochemical impact of these state variables. Concentration-based models are most often used to represent lower trophic levels in the ocean. models, forward Models, in which time is integrated forward to compute the temporal evolution of the distribution of state variables in response to a set of initial and boundary conditions. models, individually based Models, in which the biological state variables represent individual organisms or a small group of individual
103
organisms. This approach is often used when life cycles play a major role in the development of these organisms, such as is the case for most marine organisms at higher trophic levels. models, inverse Models that aim to optimally combine observations with knowledge about the workings of a system as embodied in the model. Solutions to such inverse models can be improved estimates of the current state of the system (state estimation), improved estimates of the initial or boundary conditions, or an optimal set of parameters. A typical example of an inverse model is the optimal determination of ecological parameters, such as growth and grazing rates, given, for example, the observed distribution of phytoplankton, zooplankton, and nutrients. net primary production Rate of net fixation of inorganic carbon into organic carbon by autotrophic phytoplankton. This net rate is the difference between the gross uptake of inorganic carbon during photosynthesis and autotrophic respiration. organisms, autotrophic An autotroph (from the Greek autos ¼ self and trophe ¼ nutrition) is an organism that produces organic compounds from carbon dioxide as a carbon source, using either light or reactions of inorganic chemical compounds, as a source of energy. An autotroph is known as a producer in a food chain. organisms, heterotrophic A heterotroph (Greek heterone ¼ (an)other and trophe ¼ nutrition) is an organism that requires organic substrates to get its carbon for growth and development. A heterotroph is known as a consumer in the food chain. plankton Organisms whose swimming speed is smaller than the typical speed of ocean currents, so that they cannot resist currents and are hence unable to determine their horizontal position. This is in contrast to nekton organisms that can swim against the ambient flow of the water environment and control their position (e.g., squid, fish, krill, and marine mammals). plankton, phytoplankton Phytoplankton are the (photo)autotrophic components of the plankton. Phytoplankton are pro- or eukaryotic algae that live near the water surface where there is sufficient light to support photosynthesis. Among the more important groups are the diatoms, cyanobacteria, dinoflagellates, and coccolithophorids. plankton, zooplankton Zooplankton are heterotrophic plankton that feed on other plankton. Dominant groups include small protozoans or metazoans (e.g., crustaceans and other animals).
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See also Biogeochemical Data Assimilation. Carbon Cycle. Carbon Dioxide (CO2) Cycle. Forward Problem in Numerical Models. Inverse Modeling of Tracers and Nutrients. Mesoscale Eddies. Nitrogen Cycle. Ocean Carbon System, Modeling of. Ocean Circulation. Phosphorus Cycle.
Further Reading DeYoung B, Heath M, Werner F, Chai F, Megrey B, and Monfray P (2004) Challenges of modeling ocean basin ecosystems. Science 304: 1463--1466. Doney SC (1999) Major challenges confronting marine biogeochemical modeling. Global Biogeochemical Cycles 13(3): 705--714. Fasham M (ed.) (2003) Ocean Biogeochemistry. New York: Springer.
Fasham MJR, Ducklow HW, and McKelvie SM (1990) A nitrogen-based model of plankton dynamics in the oceanic mixed layer. Journal of Marine Systems 48: 591--639. Glover DM, Jenkins WJ, and Doney SC (2006) Course No. 12.747: Modeling, Data Analysis and Numerical Techniques for Geochemistry. http://w3eos.whoi.edu/ 12.747 (accessed in March 2008). Rothstein L, Abbott M, Chassignet E, et al. (2006) Modeling ocean ecosystems: The PARADIGM Program. Oceanography 19: 16--45. Sarmiento JL and Gruber N (2006) Ocean Biogeochemical Dynamics, 526pp. Princeton, NJ: Princeton University Press.
Relevant Websites http://www.uta.edu – Interactive models of ocean ecology/biogeochemistry.
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OCEAN CARBON SYSTEM, MODELING OF S. C. Doney and D. M. Glover, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction Chemical species such as radiocarbon, chlorofluorocarbons, and tritium–3He are important tools for ocean carbon cycle research because they can be used to trace circulation pathways, estimate timescales, and determine absolute rates. Such species, often termed chemical tracers, typically have rather simple water-column geochemistry (e.g., conservative or exponential radioactive decay), and reasonably well-known time histories in the atmosphere or surface ocean. Large-scale ocean gradients of nutrients, oxygen, and dissolved inorganic carbon reflect a combination of circulation, mixing, and the production, transport, and oxidation (or remineralization) of organic matter. Tracers provide additional, often independent, information useful in separating these biogeochemical and physical processes. Biogeochemical and tracer observations are often framed in terms of ocean circulation models, ranging from simple, idealized models to full three-dimensional (3-D) simulations. Model advection and diffusion rates are typically calibrated or evaluated against transient tracer data. Idealized models are straightforward to construct and computationally inexpensive and are thus conducive to hypothesis testing and extensive exploration of parameter space. More complete and sophisticated dynamics can be incorporated into three-dimensional models, which are also more amenable for direct comparisons with field data. Models of both classes are used commonly to examine specific biogeochemical process, quantify the uptake of anthropogenic carbon, and study the carbon cycle responses to climate change. All models have potential drawbacks, however, and part of the art of numerical modeling is deciding on the appropriate model(s) for the particular question at hand. Carbon plays a unique role in the Earth’s environment, bridging the physical and biogeochemical systems. Carbon dioxide (CO2), a minor constituent in the atmosphere, is a so-called greenhouse gas that helps to modulate the planet’s climate and temperature. Given sunlight and nutrients, plants and some microorganisms convert CO2 via photosynthesis into organic carbon, serving as the building blocks and
energy source for most of the world’s biota. The concentration of CO2 in the atmosphere is affected by the net balance of photosynthesis and the reverse reaction respiration on land and in the ocean. Changes in ocean circulation and temperature can also change CO2 levels because carbon dioxide is quite soluble in sea water. In fact, the total amount of dissolved inorganic carbon (DIC) in the ocean is about 50 times larger than the atmospheric inventory. The air–sea exchange of carbon is governed by the gas transfer velocity and the surface water partial pressure of CO2 (pCO2), which increases with warmer temperatures, higher DIC, and lower alkalinity levels. The natural carbon cycle has undergone large fluctuations in the past, the most striking during glacial periods when atmospheric CO2 levels were about 30% lower than preindustrial values. The ocean must have been involved in such a large redistribution of carbon, but the exact mechanism is still not agreed upon. Human activities, including fossil-fuel burning and land-use practices such as deforestation and biomass burning, are altering the natural carbon cycle. Currently about 7.5 Pg C yr 1 (1 Pg ¼ 1015g) are emitted into the atmosphere, and direct measurements show that the atmospheric CO2 concentration is indeed growing rapidly with time. Elevated atmospheric CO2 levels are projected to heat the Earth’s surface, and the evidence for climate warming is mounting. Only about 40% of the released anthropogenic carbon remains in the atmosphere, the remainder is taken up in about equal portions (or 2 Pg C yr 1) by land and ocean sinks (Figure 1). The future magnitude of these sinks is not well known, however, and is one of the major uncertainties in climate simulations. Solving this problem is complicated because human impacts appear as relatively small perturbations on a large natural background. In the ocean, the reservoir of organic carbon locked up as living organisms, mostly plankton, is only about 3 Pg C. The marine biota in the sunlit surface ocean are quite productive though, producing roughly 50 Pg of new organic carbon per year. Most of this material is recycled near the ocean surface by zooplankton grazing or microbial consumption. A small fraction, something like 10–20% on average, is exported to the deep ocean as sinking particles or as dissolved organic matter moving with the ocean circulation. Bacteria and other organisms in the deep ocean feed on this source of organic matter from above, releasing DIC and associated nutrients back into the
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Atmosphere 750
50 ion
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Surface sediment 150 Figure 1 Schematic of global carbon cycle for the 1980s including natural background and human perturbations. Carbon inventories are in Pg C (1 Pg ¼ 1015 g) and fluxes are in Pg C yr 1. DOC, dissolved organic carbon. Adapted with permission from Schimel D, Enting IG, Heimann M, et al. (1995) CO2 and the carbon cycle. In: Houghton JT, Meira Filho LG, Bruce J, et al. (eds.) Climate Change 1994, Intergovernmental Panel on Climate Change, pp. 39–71. Cambridge, UK: Cambridge University Press.
water, a process termed respiration or remineralization. The export flux from the surface ocean is a key factor driving the marine biogeochemical cycles of carbon, oxygen, nitrogen, phosphorus, silicon, and trace metals such as iron. The surface export and subsurface remineralization of organic matter are difficult to measure directly. Biogeochemical rates, therefore, are often inferred based on the large-scale distributions of DIC, alkalinity, inorganic nutrients (nitrate, phosphate, and silicate), and dissolved oxygen. The elemental stoichiometry of marine organic matter, referred to as the Redfield ratio, is with some interesting exceptions relatively constant in the ocean, simplifying the problem of interrelating the various biogeochemical fields. Geochemical distributions have the advantage that they integrate over much of the localized time/space variability in the ocean and can be used to extrapolate to region and basin scales. Property fields, though,
reflect a combination of the net biogeochemical uptake and release as well as physical circulation and turbulent mixing. Additional information is required to separate these signals and can come from a mix of dynamical constraints, numerical models, and ocean process tracers. The latter two approaches are related because natural and artificial tracers are used to calibrate or evaluate ocean models. A key aspect of these tracers is that they provide independent information on timescale, either because they decay or are produced at some known rate, for example, due to radioactivity, or because they are released into the ocean with a known time history. The different chemical tracers can be roughly divided into two classes. Circulation tracers such as radiocarbon, tritium–3He, and the chlorofluorocarbons are not strongly impacted by biogeochemical cycling and are used primarily to quantify physical advection and mixing
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OCEAN CARBON SYSTEM, MODELING OF
rates. These tracers are the major focus here. The distribution of other tracer species is more closely governed by biology and chemistry, for example, the thorium isotope series, which is used to study export production, particle scavenging, vertical transport, and remineralization rates.
Ocean Tracers and Dynamics: A Onedimensional (1-D) Example Natural radiocarbon (14C), a radioactive isotope of carbon, is a prototypical example of a (mostly) passive ocean circulation tracer. Radiocarbon is produced by cosmic rays in the upper atmosphere and enters the surface ocean as radiolabeled carbon dioxide (14CO2) via air–sea gas exchange. The 14C DIC concentrations in the ocean decrease away from the surface, reflecting the passage of time since the water was last exposed to the atmosphere. Some radiolabeled carbon is transported to the deep ocean in sinking particulate organic matter, which can be largely corrected for in the analysis. The 14C deficits relative to the surface water can be converted into age estimates for ocean deep waters using the radioactive decay half-life (5730 years). Natural radiocarbon is most effective for describing the slow thermohaline overturning circulation of the deep ocean, which has timescales of roughly a few hundred to a thousand years. The main thermocline of the ocean, from the surface down to about 1 km or so, has more rapid ventilation timescales, from a few years to a few decades. Tracers useful in this regard are chlorofluorocarbons, tritium and its decay product 3He, and bomb radiocarbon, which along with tritium was released into the atmosphere in large quantities in the 1950s and 1960s by atmospheric nuclear weapons testing. When properly formulated, the combination of ocean process tracers and numerical models provides powerful tools for studying ocean biogeochemistry. At their most basic level, models are simply a mathematical statement quantifying the rates of the essential physical and biogeochemical processes. For example, advection–diffusion models are structured around coupled sets of differential equations: @C ¼ r ðuCÞ þ r ðKrCÞ þ J @t
½1
describing the time rate of change of a generic species C. The first and second terms on the righthand side of eqn [1] stand for the local divergence due to physical advection and turbulent mixing, respectively. All of the details of the biogeochemistry
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are hidden in the net source/sink term J, which for radiocarbon would include net input from particle remineralization (R) and radioactive loss (–l14C). One of the first applications of ocean radiocarbon data was as a constraint on the vertical diffusivity, upwelling, and oxygen consumption rates in the deep waters below the main thermocline. As illustrated in Figure 2, the oxygen and radiocarbon concentrations in the North Pacific show a minimum at mid-depth and then increase toward the ocean seabed. This reflects particle remineralization in the water column and the inflow and gradual upwelling of more recently ventilated bottom waters from the Southern Ocean. Mathematically, the vertical profiles for radiocarbon, oxygen (O2), and a conservative tracer salinity (S) can be posed as steady-state, 1-D balances: d2 S dS w 2 dz dz
½2
d2 O2 dO2 þ R O2 w dz2 dz
½3
0 ¼ Kz
0 ¼ Kz
0 ¼ Kz
d2 14C d14C þ ð14C : O2 ÞRO2 l14C w dz2 dz
½4
Kz and w are the vertical diffusivity and upwelling rates, and 14C:O2 is a conversion factor. Looking carefully at eqn [2], one sees that the solution depends on the ratio Kz/w but not Kz or w separately. Similarly the equation for oxygen gives us information on the relative rates of upwelling and remineralization. It is only by the inclusion of radiocarbon, with its independent clock due to radioactive decay, that we can solve for the absolute physical and biological rates. The solutions to eqns [2]–[4] can be derived analytically, and as shown in Figure 2 parameter values of w ¼ 2.3 10 5cm s–1, Kz ¼ 1.3 cm2 s 2, and RO2 ¼ 0:13 106 mol kg 1 yr 1 fit the data reasonably well. The 1-D model-derived vertical diffusivity is about an order of magnitude larger than estimates from deliberate tracer release experiments and microscale turbulence measurements in the upper thermocline. However, they may be consistent with recent observations of enhanced deepwater vertical mixing over regions of rough bottom topography.
Ocean Circulation and Biogeochemical Models The 1-D example shows the basic principles behind the application of tracer data to the ocean carbon
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cycle, but the complexity (if not always the sophistication) of the models and analysis has grown with time. Ocean carbon models can be roughly divided into idealized models (multi-box, 1-D and 2-D advection–diffusion models) and 3-D general circulation models (GCMs). Although the distinction can be blurry at times, idealized models are characterized typically by reduced dimensionality and/or kinematic rather than dynamic physics. That is, the circulation and mixing are specified rather than computed by the model and are often adjusted to best match the transient tracer data.
Global Box Models
An example of a simple, high-latitude outcrop box model is shown in Figures 3(a) and 3(b). The boxes represent the atmosphere, the high- and low-latitude surface ocean, and the deep interior. In the model, the ocean thermohaline circulation is represented by oneway flow with high-latitude sinking, low-latitude upwelling, and poleward surface return flow. Horizontal and vertical mixing is included by two-way exchange of water between each pair of boxes. The physical parameters are constrained so that model natural radiocarbon values roughly match observations. Note that the 14C concentration in the deep water is significantly depleted relative to the surface
boxes and amounts to a mean deep-water ventilation age of about 1150 years. The model circulation also transports phosphate, inorganic carbon, and alkalinity. Biological production, particle export, and remineralization are simulated by the uptake of these species in the surface boxes and release in the deep box. The model allows for air–sea fluxes of CO2 between the surface boxes and the atmosphere. The low-latitude nutrient concentrations are set near zero as observed. The surface nutrients in the high-latitude box are allowed to vary and are never completely depleted in the simulation of modern conditions. Similar regions of ‘high-nutrient, low-chlorophyll’ concentrations are observed in the subpolar North Pacific and Southern Oceans and are thought to be maintained by a combination of light and iron limitation as well as zooplankton grazing. The nutrient and DIC concentrations in the deep box are higher than either of the surface boxes, reflecting the remineralization of sinking organic particles. The three-box ocean outcrop model predicts that atmospheric CO2 is controlled primarily by the degree of nutrient utilization in high-latitude surface regions. Marine production and remineralization occur with approximately fixed carbon-to-nutrient ratios, the elevated nutrients in the deep box are associated with an equivalent increase in DIC and pCO2. Large adjustments in the partitioning of
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OCEAN CARBON SYSTEM, MODELING OF
(a)
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Figure 3 Results from a simple three-box ocean carbon cycle model. (a) The physical circulation and modeled radiocarbon (D14C) values. (b) The model biogeochemical fields, ocean DIC, and phosphate (PO4) and atmospheric pCO2. From Toggweiler JR and Sarmiento JL, Glacial to inter-glacial changes in atmospheric carbon dioxide: The critical role of ocean surface waters in high latitudes, The Carbon Cycle and Atmospheric CO2: Natural Variations Archean to Present, Sundquist ET and Broecker WS (eds.), pp. 163–184, 1985, Copyright [1985]. American Geophysical Union. Adapted by permission of American Geophysical Union.
carbon between the ocean and atmosphere can occur only where this close coupling of the carbon and nutrient cycles breaks down. When subsurface water is brought to the surface at low latitude, production draws the nutrients down to near zero and removes to first order all of the excess seawater DIC and pCO2. Modifications in the upward nutrient flux to the low latitudes have relatively little impact on the
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model atmospheric CO2 as long as the surface nutrient concentrations stay near zero. At high latitudes, however, the nutrients and excess DIC are only partially utilized, resulting in higher surface water pCO2 and, over decades to centuries, higher atmospheric CO2 concentrations. Depending on the polar biological efficiency, the model atmosphere effectively sees more or less of the high DIC concentrations (and pCO2 levels) of the deep ocean. Thus changes in ocean biology and physics can have a correspondingly large impact on atmospheric CO2. On longer timescales (approaching a few millennia), these variations are damped to some extent by adjustments of the marine calcium carbonate cycle and ocean alkalinity. The three-box outcrop model is a rather crude representation of the ocean, and a series of geographical refinements have been pursued. Additional boxes can be added to differentiate the individual ocean basins (e.g., Atlantic, Pacific, and Indian), regions (e.g., Tropics and subtropics), and depths (e.g., thermocline, intermediate, deep, and bottom waters), leading to a class of models with a halfdozen to a few dozen boxes. The larger number of unknown advective flows and turbulent exchange parameters, however, complicates the tuning procedure. Other model designs take advantage of the vertical structure in the tracer and biogeochemical profile data. The deep box (es) is discretized in the vertical, essentially creating a continuous interior akin to a 1-D advection–diffusion model. This type of model was often used in the 1970s and 1980s for the initial anthropogenic CO2 uptake calculations, where it is important to differentiate between the decadal ventilation timescales of the thermocline and the centennial timescales of the deep water. Intermediate Complexity and Inverse Models
In terms of global models, the next step up in sophistication from box models is intermediate complexity models. These models typically have higher resolution and/or include more physical dynamics but fall well short of being full GCMs. Perhaps the most common examples for ocean carbon cycle research are zonal average basin models. The dynamical equations are similar to a GCM but are integrated in 2-D rather than in 3-D, the third east–west dimension removed by averaging zonally across the basin. In some versions, multiple basins are connected by an east–west Southern Ocean channel. The zonal average models often have a fair representation of the shallow wind-driven Ekman and deep thermohaline overturning circulations but obviously lack western boundary currents and gyre circulations. Tracer data
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remain an important element in tuning some of the mixing coefficients and surface boundary conditions and in evaluating the model solutions. Based on resolution, many inverse models can also be categorized as intermediate complexity, but their mode of operation differs considerably from the models considered so far. In an inverse model, the circulation field and biogeochemical net source/sink (the J terms in the notation above) are solved for using the observed large-scale hydrographic and tracer distributions as constraints. Additional dynamic information may also be incorporated such as the geostrophic velocity field, general water mass properties, or float and mooring velocities. Inverse calculations are typically posed as a large set of simultaneous linear equations, which are then solved using standard linear algebra methods. The inverse techniques are most commonly applied to steady-state tracers, though some exploration of transient tracers has been carried out. The beauty of the inverse approach is that it tries to produce dynamically consistent physical and/or biogeochemical solutions that match the data within some assigned error. The solutions are often underdetermined in practice, however, which indicates the existence of a range of possible solutions. From a biogeochemical perspective, the inverse circulation models provide estimates of the net source/sink patterns, which can then be related to potential mechanisms. Thermocline Models
Ocean process tracers and idealized models have also been used extensively to study the ventilation of the main thermocline. The main thermocline includes the upper 1 km of the tropical to subpolar ocean where the temperature and potential density vertical gradients are particularly steep. Thermocline 5000
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ventilation refers to the downward transport of surface water recently exposed to the atmosphere, replenishing the oxygen and other properties of the subsurface interior. Based on the vertical profiles of tritium and 3He as well as simple 1-D and box models, researchers in the early 1980s showed that ventilation of the main thermocline in the subtropical gyres occurs predominately as a horizontal process along surfaces of constant density rather than by local vertical mixing. Later work on basinscale bomb-tritium distributions confirmed this result and suggested the total magnitude of subtropical ventilation is large, comparable to the total winddriven gyre circulation. Two dimensional gyre-scale tracer models have been fruitfully applied to observed isopycnal tritium–3He and chlorofluorocarbon patterns. As shown in Figure 4, recently ventilated water (nearzero tracer age) enters the thermocline on the poleward side of the gyre and is swept around the clockwise circulation of the gyre (for a Northern Hemisphere case). Comparisons of model tracer patterns and property–property relationships constrain the absolute ventilation rate and the relative effects of isopycnal advection versus turbulent mixing by depth and region. Thermocline tracer observations are also used to estimate water parcel age, from which biogeochemical rate can be derived. For example, remineralization produces an apparent oxygen deficit relative to atmospheric solubility. Combining the oxygen deficit with an age estimate, one can compute the rate of oxygen utilization. Similar geochemical approaches have been or can be applied to a host of problems: nitrogen fixation, denitrification, dissolved organic matter remineralization, and nutrient resupply to the upper ocean. The biogeochemical application of 2-D gyre models has not been pursued in as much detail.
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Figure 4 Tracer results from a two-dimensional gyre model. The model represents the circulation on a constant density surface (isopycnal) in the main thermocline that outcrops along the northern boundary (shaded region). Thermocline ventilation is indicated by the gradual increase of tritium–3He ages around the clockwise flowing gyre circulation. TU, tritium unit (1 TU ¼ 1 3H atom/1018 H atoms). Reproduced with permission from Musgrave DL (1990) Numerical studies of tritium and helium-3 in the thermocline. Journal of Physical Oceanography 20: 344–373.
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This form of modeling, however, is particularly useful for describing regional patterns and in the areas where simple tracer age approaches breakdown.
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By their very nature, idealized models neglect many important aspects of ocean dynamics. The alternative is full 3-D ocean GCMs, which incorporate more realistic spatial and temporal geometry and a much fuller suite of physics. There are several different families of ocean physical models characterized by the underlying governing equations and the vertical discretization schemes. Within model families, individual simulations will differ in important factors of surface forcing and choice of physical parametrizations. Often these parametrizations account for complex processes, such as turbulent mixing, that occur on small space scales that are not directly resolved or computed by the model but which can have important impacts on the larger-scale ocean circulation; often an exact description of specific events is not required and a statistical representation of these subgrid scale processes is sufficient. For example, turbulent mixing is commonly treated using equations analogous to those for Fickian molecular diffusion. Ocean GCMs, particularly coarse-resolution global versions, are sensitive to the subgrid scale parametrizations used to account for unresolved processes such as mesoscale eddy mixing, surface and bottom boundary layer dynamics, and air–sea and ocean–ice interactions. Different models will be better or worse for different biogeochemical problems. Considerable progress has been achieved over the last two decades on the incorporation of chemical tracers, biogeochemistry, and ecosystem dynamics into both regional and global 3-D models. Modeling groups now routinely simulate the more commonly measured tracers (e.g., radiocarbon, chlorofluorocarbons, and tritium–3He). Most of these tracers have surface sources and, therefore, provide clear indications of the pathways and timescales of subsurface ventilation in the model simulations. The bottom panel of Figure 5 shows from a simulation of natural radiocarbon the layered structure of the intermediate and deep water flows in the Atlantic basin, with Antarctic Intermediate and Bottom Waters penetrating from the south and southwardflowing North Atlantic Deep Water sandwiched in between. In a similar fashion, biogeochemical modules have been incorporated to simulate the 3-D fields of nutrients, oxygen, DIC, etc. Just as with the physics, the
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degree of biological complexity can vary considerably across ocean carbon models. At one end are diagnostic models where the production of organic matter in the surface ocean is prescribed based on satellite ocean color data or computed implicitly by forcing the simulated surface nutrient field to match observations (assuming that any net reduction in nutrients is driven by organic matter production). At the other
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Circumpolar Deep Water Δ14C (per mil) Figure 6 Transient tracer constraints on numerical model ocean carbon dynamics. Each panel shows results for the suite of global ocean simulations that participated in the international Ocean Carbon-Cycle Model Intercomparison Project (OCMIP). The points with the error bars are the corresponding observations. The x-axis in all panels is the average Circumpolar Deep Water radiocarbon level (D14C, per mil). The y-axis in the left column is the chlorofluorocarbon (CFC-11) inventory (108 mol) outside of the Southern Ocean (top row) and in the Southern Ocean. The y-axis in the right column is the corresponding anthropogenic carbon inventory (Pg C). From Matsumoto K, Sarmiento JL, Key RM, et al., Evaluation of ocean carbon cycle models with data-based metrics, Geophysical Research Letters, Vol. 31, L07303 (doi:10.1029/2003GL018970), 2004. Copyright [2004] American Geophysical Union. Reproduced by permission of American Geophysical Union.
extreme are simulations that couple prognostic treatments of multiple phytoplankton, zooplankton, and bacteria species to the biogeochemical tracers. Transient tracers enter the picture because the biogeochemical distributions are also strongly influenced circulation. This is nicely illustrated by comparing the simulated radiocarbon and DIC sections in Figure 5. To first order, older waters marked by low radiocarbon exhibit high DIC levels due to the accumulation of carbon released from particle remineralization. Tracers can be used to guide the improvement of the physical circulation and highlight areas where the circulation is poor and thus the biogeochemical tracers may be suspect. They can also
be used more quantitatively to create model skill metrics; this approach can be applied across multiple model simulations to provide the best overall constraint on quantities like the ocean uptake of anthropogenic CO2 or export production (Figure 6).
Models as Research Tools Designed around a particular question or hypothesis, conceptual models attempt to capture the basic elements of the problem while remaining amenable to straightforward analysis and interpretation. They are easy to construct and computationally
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inexpensive, requiring only a desktop PC rather than a supercomputer. When well-formulated, idealized models provide a practical method to analyze ocean physical and biogeochemical dynamics and in some cases to quantitatively constrain specific rates. Their application is closely tied to ocean tracer observations, which are generally required for physical calibration and evaluation. Idealized models remain a valuable tool for estimating the oceanic uptake of anthropogenic carbon and the long time-scale responses (centuries to millennia) of the natural carbon cycle. Also, some of the more memorable and lasting advances in tracer oceanography are directly linked to simple conceptual models. Examples include constraints on the deep-water large-scale vertical diffusivity and demonstration of the dominance of lateral over vertical ventilation of the subtropical main thermocline. Three-dimensional models offer more realism, at least apparently, but with the cost of greater complexity, a more limited number of simulations, and a higher probability of crucial regional errors in the base solutions, which may compromise direct, quantitative model-data comparisons. Ocean GCM solutions, however, should be exploited to address exactly those problems that are intractable for simpler conceptual and reduced dimensional models. For example, two key assumptions of the 1-D advection–diffusion model presented in Figure 2 are that the upwelling occurs uniformly in the horizontal and vertical and that mid-depth horizontal advection is not significant. Ocean GCMs and tracer data, by contrast, show a rich three-dimensional circulation pattern in the deep Pacific. The behavior of idealized models and GCMs can diverge, and it is not always clear that complexity necessarily leads to more accurate results. In the end, the choice of which model to use depends on the scientific problem and the judgment of the researcher. Probably the best advice is to explore solutions from a hierarchy of models and to thoroughly evaluate the skill of the models against a range of tracers and other dynamical measures. Just because a model does a good job reproducing the distribution of one tracer field does not mean that it can be applied indiscriminately to another variable, especially if the underlying dynamics or timescales differ. Models can be quite alluring in the sense that they provide concrete answers to questions that are often difficult or nearly impossible to address from sparsely sampled field data. However, one should not forget that numerical models are simply a set of tools for doing science. They are no better than the foundations upon which they are built and should not be carried out in isolation from observations of the real
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ocean. For ocean carbon cycle models, the two key elements are the ocean physical circulation and the biogeochemical processes. Even the best biogeochemical model will perform poorly in an illconstructed physical model. Conversely, if the underlying biogeochemical mechanisms are poorly known, a model may be able to correctly reproduce the distributions of biogeochemical tracers but for the wrong reasons. Mechanistic-based models are critical in order to understand and predict natural variability and the response of ocean biogeochemistry to perturbations such as climate change.
Glossary anthropogenic carbon The additional carbon that has been released to the environment over the last several centuries by human activities including fossil-fuel combustion, agriculture, forestry, and biomass burning. excess 3He Computed as the 3He in excess of gas solubility equilibrium with the atmosphere. export production That part of the organic matter formed in the surface layer by photosynthesis that is transported out of the surface layer and into the interior of the ocean by particle sinking, mixing, and circulation, or active transport by organisms. radiocarbon 14C Either d14C or D14C where d14 C ¼ [(14C/12C)sample/(14C/12C)standard – 1] 1000 (in parts per thousand or ‘per mil’); D14C is similar but corrects the sample 14C for biological fractionation using the sample 13C/12C ratio. transient tracers Chemical tracers that contain time information either because they are radioactive or because their source, usually anthropogenic, has evolved with time. tritium–3He age An age is computed assuming that all of the excess 3He in a sample is due to the radioactive decay of tritium, age ¼ ln[(3H þ 3He)/ 3 H ]/l, where l is the decay constant for tritium. ventilation The physical process by which surface properties are transported into the ocean interior.
See also Biogeochemical Data Assimilation. Carbon Cycle. Carbon Dioxide (CO2) Cycle. CFCs in the Ocean. Dispersion and Diffusion in the Deep Ocean. Forward Problem in Numerical Models. Nitrogen Cycle. Ocean Circulation. Ocean Circulation: Meridional Overturning Circulation. Radiocarbon. Tritium–Helium Dating.
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Further Reading Broecker WS and Peng T-H (1982) Tracers in the Sea. Palisades, NY: Lamont-Doherty Geological Observatory, Columbia University. Charnock H, Lovelock JE, Liss P, and Whitfield M (eds.) (1988) Tracers in the Ocean. London: The Royal Society. Doney SC, Lindsay K, Caldeira K, et al. (2004) Evaluating global ocean carbon models: The importance of realistic physics. Global Biogeochemical Cycles 18: GB3017 (doi:10.1029/2003GB002150). England MH and Maier-Reimer E (2001) Using chemical tracers in ocean models. Reviews in Geophysics 39: 29--70. Fasham M (ed.) (2003) Ocean Biogeochemistry. New York: Springer. Jenkins WJ (1980) Tritium and 3He in the Sargasso Sea. Journal of Marine Research 38: 533--569. Kasibhatla P, Heimann M, and Rayner P (eds.) (2000) Inverse Methods in Global Biogeochemical Cycles. Washington, DC: American Geophysical Union. Matsumoto K, Sarmiento JL, Key RM, et al. (2004) Evaluation of ocean carbon cycle models with databased metrics. Geophysical Research Letters 31: L07303 (doi:10.1029/2003GL018970). Munk WH (1966) Abyssal recipes. Deep-Sea Research 13: 707--730. Musgrave DL (1990) Numerical studies of tritium and helium-3 in the thermocline. Journal of Physical Oceanography 20: 344--373. Sarmiento JL and Gruber N (2006) Ocean Biogeochemical Dynamics. Princeton, NJ: Princeton University Press.
Schimel D, Enting IG, Heimann M, et al. (1995) CO2 and the carbon cycle. In: Houghton JT, Meira Filho LG, and Bruce J, et al. (eds.) Climate Change 1994, Intergovernmental Panel on Climate Change, pp. 39--71. Cambridge, UK: Cambridge University Press. Siedler G, Church J, and Gould J (eds.) (2001) Ocean Circulation and Climate. New York: Academic Press. Siegenthaler U and Oeschger H (1978) Predicting future atmospheric carbon dioxide levels. Science 199: 388--395. Toggweiler JR and Sarmiento JL (1985) Glacial to interglacial changes in atmospheric carbon dioxide: The critical role of ocean surface waters in high latitudes. In: Sundquist ET and Broecker WS (eds.) The Carbon Cycle and Atmospheric CO2: Natural Variations Archean to Present, pp. 163--184. Washington, DC: American Geophysical Union.
Relevant Websites http://cdiac.ornl.gov – Global Ocean Data Analysis Project, Carbon Dioxide Information Analysis Center. http://w3eos.whoi.edu – Modeling, Data Analysis, and Numerical Techniques for Geochemistry (resource page for MIT/WHOI course number 12.747), at Woods Hole Oceanographic Institution. http://us-osb.org – Ocean Carbon and Biogeochemistry (OCB) Program http://www.ipsl.jussieu.fr – Ocean Carbon-Cycle Model Intercomparison Project (OCMIP), Institut Pierre Simon Laplace.
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OCEAN CIRCULATION N. C. Wells, Southampton Oceanography Centre, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1936–1946, & 2001, Elsevier Ltd.
Introduction This article discusses the following aspects of ocean circulation: what is meant by the term ocean circulation; how the ocean circulation is determined by measurements and dynamical processes; the consequences of this circulation on the Earth’s climate.
and satellites (see Satellite Altimetry), combined with the mathematical methods of dynamical oceanography (see Elemental Distribution: Overview. General Circulation Models. Inverse Models). This integrated approach allows hypotheses to be made that can be tested by comparison with observations. Furthermore, mathematical models of the ocean circulation, based on the dynamical principles, can be constructed and tested against observations (see El Nin˜o Southern Oscillation (ENSO) Models. Regional and Shelf Sea Models). This article considers how ocean circulation is measured, how the major processes at work are determined and the consequences of the ocean circulation on the climate system.
What is Ocean Circulation? The ocean circulation in its simplest form is the movement of sea water through the ocean, which principally transfers temperature and salinity, from one region to another. Temperature differences between regions cause heat transfers. Similarly, differences in salinity produce transfers of salt. On the time scale of the ocean circulation the inputs and exports of salt into and out of the ocean make a negligible contribution to overall salinity and so variations in salinity occur by the addition and removal of fresh water into and out of the ocean. Two major processes control the ocean circulation: the action of the wind and the action of small density differences, produced by differences in temperature and salinity, within the ocean. The former process is the wind driven circulation (see Wind Driven Circulation) whereas the latter is the thermohaline circulation. Although it is useful to separate these two processes to better understand the ocean circulation, they are not independent from each other. Ocean circulation is in reality a very complex system, as the flows are not steady in time or space. They are turbulent flows that show variability on scales from the largest scale of the ocean basins (see El Nin˜o Southern Oscillation (ENSO), North Atlantic Oscillation (NAO)) to the smallest scales where the energy is finally dissipated as heat (seeWaves on Beaches). This turbulent structure of the ocean means there are fundamental limitations on the predictability of its behavior. Because of this inherent complexity oceanographers have approached ocean circulation by using a combination of observational methods including ships (see Ships), buoys (see Drifters and Floats) (see Moorings)
How is the Ocean Circulation Determined? The determination of ocean currents involves measurement of the displacement of an element of fluid over a measured time interval. The position of the measurement is defined mathematically in a Cartesian coordinate system (Figure 1) where x is positive eastward direction (lines of constant latitude), y is positive northward direction (geographic North Pole), and z is positive upwards; z ¼ 0 corresponds to mean sea level. Without ocean currents and tides the sea level would be an equipotential surface, i.e., one of constant potential energy. The z coordinate is perpendicular to the equipotential surface. The origin is the intersection of the Greenwich meridian (Universal meridian) and the equator with mean sea level. The coordinates of a parcel of water can be determined by the Global Positioning System (GPS). This satellite-based system provides a horizontal position with an accuracy of better than 100 m, which is sufficient for large-scale flows in the ocean. Large-scale flows are at least 10 km in spatial scale and have timescales of at least a day. A pressure device attached to the current meter normally determines the vertical position. There are two mathematical methods for defining the displacement of the fluid. One is to measure the velocity of the fluid at a fixed point in the ocean, and the other is to follow the element of the fluid and to measure its velocity as it moves through the ocean. The first method is known as a Eulerian description and the second is a Lagrangian description of flow. In
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Eulerian
v (x, y, z, t ) Measurement of velocity of the fluid at fixed point (x, y, z )
z
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Having defined the two mathematical methods how the currents are measured in practice is now considered. Initially, these methods will address only the measurement of the horizontal flow. The vertical flow is difficult to measure directly, and will be discussed later in this article. First, the Eulerian method is considered. The measurement of the flow at a fixed point in the ocean is only straightforward when a fixed position can be maintained, for instance with a current meter attached to the bottom of the ocean or to a pier on the coast. Most measurements have to be made well away from land. This is achieved by attaching the current meter to a mooring (see Moorings) which is attached to weights and then deployed (Figure 2). . The position of a mooring can be determined by GPS. The current meter may be a rotary device or an acoustic device. The rotary current meter measures the number of revolutions over a fixed period, whereas the acoustic one measures the change in frequency of an emitted sound pulse caused by the ocean current (i.e., it uses the Doppler effect). Moorings may be deployed for periods of up to 2 years. In the analysis of the record it is normal to
v (a,t ) Measurement of the velocity of a fluid element. a is the position vector from the origin to the fluid element.
Reflector Tripod tower
P′
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Element of fluid moves a small distance a in a time t, from P to P′.
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The velocity is the a t In infinitesimal time a = v (a,t ) t Anchor weight
Acoustic release for recovery
Figure 1 Eulerian and Lagrangian specifications of flow.
principle, the two methods are independent of each other. This means that a Eulerian measurement can not provide Lagrangian currents, and Lagrangian measurements can not provide Eulerian currents.
Figure 2 A typical current meter mooring.
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float would use an acoustic navigation system. Some floats can descend to a predetermined depth, maintain that depth for a few weeks and then return to the surface for a position fix. This technique allows the current to be measured down to depths of 1 km below the surface. Each method gives different information on the flow field. A mooring will give a time series of the horizontal current, whereas a float will give the trajectory of the horizontal displacement of the parcel. It is worth remarking that most of the information on the surface ocean circulation has come from mariners’ observations of the ships set, a method which has been used since the nineteenth century. However, these measurements have their limitations since they are neither eulerian nor lagrangian measurements and additional influences (e.g., wind effects on the ship) may cause errors. This information can be analyzed in many different ways to discern the major current systems. From a set of moorings deployed for a few years across, say, the Gulf Stream, the mean flow (i.e., the average of all the current measurements) and its variability can be determined. The mean flow could be calculated over a particular time period. This time period is limited by the period of deployment, which is of the order of 2 years. This is rather short for a climatological mean, and a much longer period of 10 years is desirable. A few longer time series of currents have been determined for the Gulf Stream in the Florida Straits (see Rigs and offshore Structures) and for the Antarctic Circumpolar Current in the Drake Passage (see Antarctic Circumpolar Current). Recall ocean currents are turbulent and therefore have variability on a whole range of timescales. Hence the mean flow gives no information on the variability of the flow. However, the statistics of the flow can be calculated, based on the kinetic energy (KE). The kinetic energy/unit volume is defined as:
3_ 6 m aluminum mast with flag, light, and radar reflector
Styrofoam float
3_ 6 m iron pipe
Chain ballast
Manila line
1.3 mm piano wire
8.5 m parachute
Manila line
Chain ballast
Figure 3 A typical drifter with a parachute drogue of a few meters below the surface. It will follow the current at the depth of the parachute.
remove the high frequency variability of less than 1 day caused by tides by filtering the data. A Lagrangian measurement of current can be determined by following an element of water with a float. The horizontal displacement of the water over a small interval of time defines the Lagrangian current. Figure 3 shows typical float designs that are used. (see Drifters and Floats). The position of the float can be determined by two methods. A float that has a surface satellite transmitter/receiver can have its position determined by GPS, whereas a subsurface
1 KE ¼ r u2 þ v2 2 where r is the density of the sea water and u and v are the eastward and northward components of the horizontal flow, respectively. If the time mean current is defined as u¯ and u0 as the deviation from u¯ at any time, the mean kinetic energy (KEM) and eddy kinetic energy (EKE) can be defined by:
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1 KEM ¼ r u¯ 2 þ v¯ 2 2 1 EKE ¼ r u02 þ v02 2
OCEAN CIRCULATION
_ 0.5 _ 1.0
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intense and highly variable current off the coast of South Africa. Although this ratio gives a measure of the variability of the current, it does not give any idea of the exact time or space scales over which the current is varying. For example, the current may show a slow change from one season to another or it may show faster variation due to eddies. To address this variation time series analysis, such as Fourier analysis, can be used to determine the KE of the flow for different time periods. Fourier analysis produces a spectrum of the KE, either in frequency for a time series, or in wave number for a spatial variation in flow. Figure 5 shows the analysis of a time series into its component frequencies. If the current is varying on all timescales the spectrum would be flat, but if there was only one dominant period, it would peak at that one frequency. This particular analysis shows that the current is varying at the tidal frequency and the inertial frequency both
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_ 5.0
_ 42 _ 41 _ 40 _ 39 _ 38 _ 37 _ 36 _ 35 _ 34 _ 33 Latitude
Figure 4 The mean kinetic energy (KEM) (A) and the eddy kinetic energy (EKE) (B) in a north–south slice through the Agulhas Current system at 14.41E. The KEM maximum corresponds to the mean position of the Agulhas Return Current (Eastward flow) between 401 and 411S, and the Agulhas Current (Westward flow) between 371S and 381S. The EKE distribution is much broader than KEM, which shows the large horizontal extent of the flow variability. The ratio of EKE/ KEM is typically about a third, which indicates a very variable current system. (Reproduced from Wells NC, Ivchenko V and Best SE (2000) Instabilities in the Agulhas Retroflection Current system: A comparative model study. Journal of Geophysical Research 105: 3233–3246.)
These two numbers give quantitative measures on the mean and variability of the flow respectively. The ratio EKE/KEM gives a measure of the relative variability of the flow. If the ratio is very much less than 1 then the flow is steady, whereas if the ratio is approximately equal to 1 then the flow is very variable. Figure 4 shows the variability of the flow in the Agulhas Current (see Agulhas Current), which is an
Inertia Semidiurnal tide
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Period (h) Figure 5 Frequency spectrum of kinetic energy from four depths at site D (391N, 701W), north of the Gulf Stream. Note the two high-frequency peaks, coinciding with the inertial period (19 h) and the semidiurnal tide (12.4 h). (Reproduced from Rhines PB (1971) A note on long-period motions at site D. Deep Sea Research 18: 21–26.)
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at the high frequency end of the spectrum. The inertial frequency is given by 2O sin f where O is the rotation rate of the earth and f is the latitude. At the lower frequency end, which corresponds to the ocean circulation frequency, there is a broad band of high kinetic energy. This band is due to eddies (see Mesoscale Eddies) which cause fluctuations of currents on timescales of weeks to months. For these mean climatological currents, our knowledge has been augmented by the application of the dynamic method. This method is based on the observation that large-scale ocean currents are in geostrophic balance, over large areas of the ocean. Geostrophic balance means that the Coriolis force balances the horizontal pressure gradient force. The geostrophic flow is a good approximation to the flow in the interior of an ocean outside the equatorial region. The horizontal pressure gradient is dependent on the slope of the ocean surface and the horizontal variation of the density distribution within the ocean. In the future, the former may be determined by satellite measurements of the sea surface height and the geoid1 but at present we do not have an accurate geoid at sufficiently high resolution to measure the sea surface slope. The latter can be determined from temperature, salinity and pressure measurements that have been made over large ocean areas during the last century. The dynamic method allows the determination of the vertical shear of the horizontal geostrophic current, and therefore to determine the absolute geostrophic current, additional measurements are required. For example if the current has been measured at a particular depth then the dynamic method can be referenced to that depth and the vertical profile of current can be obtained. The recent World Ocean Circulation Experiment (WOCE) hydrographic program has provided more measurements of the ocean than all previous hydrographic programs and will give the most comprehensive assessment of climatological horizontal ocean flow to date. Recall that the vertical circulation of the ocean cannot be measured directly because it is technically too difficult. Current indirect methods used to determine the vertical circulation rely on the use of mathematical approaches, such as dynamical models, or the use of chemical tracers. Observations of temperature and salinity can be inserted into a mathematical ocean general circulation model (see below and Elemental Distribution:
1
The geoid is an equipotential surface, which would be represented by the sea level of a stationary ocean. Ocean currents cause deviations in sea level from the geoid.
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Overview) which allows the three-dimensional, circulation to be determined, subject to limitations in the accuracy of the model.. Chemical tracers have been inadvertently injected into the ocean from nuclear tests in the 1960s and from industrial processes (e.g., chlorofluorocarbons). (see CFCs in the Ocean. Naturally occurring tracers such as 14C also exist. These tracers can be measured with high accuracy in a few laboratories around the world and from their distributions at different times, the three-dimensional circulation can be estimated. This method reveals the time history of the ocean circulation wherever the tracer is measured. (see Tritium–Helium Dating). This is very different information from that provided by the methods previously discussed, but nonetheless it can reveal unique aspects of the flow. For example, nuclear fallout deposited in the surface layer of the Nordic seas in the 1960s was located in the deep western boundary current 10 years later.
An Ocean General Circulation Model An ocean general circulation model is composed of a set of mathematical equations which describe the time-dependent dynamical flows in an ocean basin. The basin is discretized into a set of boxes of regular horizontal dimensions but variable thickness in the vertical dimension. The horizontal flow (northward and eastward components) is predicted by the momentum equation (Figure 6A) at the corners of each box (Figure 7). The forcing for the flow may come from the wind stress (the frictional term in the momentum equation) and from the surface buoyancy fluxes, arising from heat and freshwater (precipitation þ runoff– evaporation) exchange with the atmosphere and adjacent landmasses. These buoyancy fluxes change the temperature and salinity in the surface layer of the ocean. The surface water masses are then subducted into the interior of the ocean by the vertical and horizontal components of the flow, where they are mixed with other water masses. The processes of transport and mixing are described by the temperature and salinity equations (Figure 6B and C), at the center of each ocean box (Figure 7). From these two equations the seawater density, and thence the pressure can be obtained for each box. The horizontal pressure gradient is then determined for the momentum equation, and the vertical velocity is calculated from the horizontal divergence of the flow. This set of time-dependent equations can then be used to describe all the dynamical components of the flow field, provided that suitable initial and boundary conditions are specified.
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Momentum equation Rate of change = of horizontal momentum in box
Transport of momentum + across faces of box
Coriolis + acceleration
(Rotation of earth)
(A)
Horizontal pressure gradient
+
Friction
(Internal dissipation (Gradients in sea level and in density) and at ocean boundaries)
Temperature equation Rate of change = of temperature in box
Transport of temperature + across faces of box
Mixing of temperature between box and adjacent boxes
(B)
Salinity equation Rate of change = of salinity in box
(C)
Transport of salinity across faces of box
+
Mixing of salinity between box and adjacent boxes
Figure 6 The basic equations for an ocean general circulation model. (A) Momentum equation; (B) temperature equation; (C) salinity equation. (Reproduced from Summerhayes and Thorpe, 1996.)
Heat loss
Earth′s rotation
Wind
u
u
T u Su
Figure 7 A schematic of the model boxes in an ocean general circulation model. The equations (Figure 6) for momentum are solved at the corners of the boxes (u), while the temperature (T), and salinity (S) equations are solved at the centers of boxes. The model is forced by climatological wind stress, surface heat fluxes, and freshwater fluxes. (Reproduced from Summerhayes and Thorpe, 1996.)
Wind-driven and Thermohaline Circulation The wind-driven circulation (see Wind Driven Circulation) is considered first. The surface layer of the ocean is directly driven by the surface wind stress
and is also subject to the exchange of heat and fresh water between ocean and atmosphere. This layer, which is typically less than 100 m in depth, is referred to as the Ekman layer. That is a steady wind stress causes a transport of the surface water 901 to the right of the wind direction in the Northern Hemisphere and 901 to the left in the Southern Hemisphere. This is due to the combined action of the wind stress on the ocean surface and the Coriolis force. These Ekman flows can converge and produce a downwelling flow into the interior of the ocean. Conversely a divergent Ekman transport will produce an upwelling flow from the interior into the surface layer. This type of flow is known as Ekman pumping, and is directly related to the Curl of the Wind Stress (see Wind Driven Circulation). It is of fundamental importance for the driving of the large-scale horizontal circulation, in the upper layer of the ocean. For example, between 301 and 501 latitude the climatological westerly wind, drives an Ekman flow equatorward, whereas between 151 and 301 latitude the trade winds drive an Ekman flow polewards. At about 301 latitude the flows converge and sink into the deeper ocean. Before discussing the influence of Ekman pumping on the interior ocean circulation the role of density is considered. The density of sea water increases with depth. From hydrographic measurements of density, the horizontal variation of the depth of a chosen density surface can be mapped. These constant density surfaces are known as isopycnals. The flow tends to move along these surfaces and therefore the variations in the depth of these surfaces gives a picture of
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the horizontal flow in the deep ocean, away from the surface layer and benthic layer. The isopycnal surfaces dip down in the center of the subtropical gyre at about 301. The formation of this lens of light warm water is related to the climatological distribution of surface winds, which produce a convergence of Ekman transport towards the center of the gyre, and a downwelling of surface waters into the interior of the ocean. At the center of the lens, the sea surface domes upwards reaching a height of 1 m above the sea surface at the rim. Due to hydrostatic forces the main thermocline is depressed downwards to depths of the order of 500–1000 m (Figure 8). The surface horizontal circulation flows anticyclonically around the lens with the strongest currents on the western edge, where the slope of the density surface reaches a maximum. These are geostrophic currents, where there is a balance between
2 km H
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the Coriolis force and the horizontal pressure gradient force. Generally, the circulation in the subtropical gyres is clockwise in the Northern Hemisphere and anticlockwise in the Southern Hemisphere. These large-scale horizontal gyres are ultimately caused by the climatological surface wind circulation and are found in all the ocean basins. The surface layer is also subject to heating and cooling, and the exchange of fresh water between ocean and atmosphere, both of which will change the density of the layer. For example, heat is lost over the Gulf Stream on the rim of the light water lens of the subtropical gyre. Recall that flow tends to follow isopycnal layers and these layers will slope downwards towards the center of the gyre. Cooling of the waters in the Gulf Stream leads to the sinking of surface waters to produce a water mass known as 181C water. This water, which is removed from the surface layer, will slowly move along the isopycnal layers into the thermocline. As it moves clockwise around the gyre it will be subducted in to the deeper layers of the thermocline, in a spiral-like motion (Figure 9). The deepest extent of the main thermocline is located in the subtropical gyre to the west of
Cooling
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Upwelling in the subpolar gyre
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Figure 8 A representation of the meridional average section through the atmosphere for December–February in the Northern Hemisphere. The cells are the Hadley Cell (H) and Ferrel Cell (F). The strength of the cells is represented by the solid contours which are in units of 40 megatonnes/second, whilst the dashed contours are in units of 20 megatonnes/second. Note the predominantly downward motion at B30 degree latitude, associated with the subtropical anticyclones, and the strong upward motion at equatorial latitudes which is associated with the Inter-Tropical Convergence Zone. A meridional transect through the Atlantic Ocean, showing the position of the main thermocline. The small arrows represent the wind driven downwelling (Ekman pumping) at B30 degree latitude, and the equatorial upwelling, which occurs within and above the main thermocline. The North Atlantic Deep Water (NADW) is produced in the Labrador and Nordic Seas and is the predominant deep water mass by volume. The Antarctic Intermediate Water (AIW) is produced at B501S, and by virtue of its salinity is lighter than the NADW. In contrast Antarctic Bottom Water is the most dense water mass in the worlds ocean and is formed in the Weddell and Ross Seas. These deep flows form part of the thermo-haline circulation. The vertical scales are exaggerated in the lower troposphere and in the upper ocean. The horizontal scale is proportional to the area of the Earth’s surface between latitude circles.
Heat and salt diffuse across pycnocline
Pyc
noc
Subpolar gyre
line
Subtropical gyre Below the pycnocline the deep waters have more uniform temperature and salinity characteristics
Figure 9 Schematic representation of the wind-driven circulation in the subpolar and subtropical gyre of an ocean basin. The wind circulation causes a convergence of Ekman transport to the center of the subtropical gyre and downwelling into the interior. Conversely in the subpolar gyre there is a divergence of the Ekman transport and upwelling from the interior into the surface layer. This Ekman pumping is responsible for the gyre circulations (see text for details). The western boundary currents are depicted by the closeness of the streamlines. They are caused by the poleward change in the Coriolis force known as the BETA effect. (Reproduced from Bean MS (1997) PhD thesis, University of Southampton.)
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Bermuda on the eastern edge of the Gulf Stream rather than in the center of the ocean basin. This asymmetry of the gyre is related to the beta effect, i.e., the change of the Coriolis parameter with latitude (see Agulhas Current). The subtropical gyres are one of the most wellstudied regions of the ocean, and our understanding is therefore most developed in these regions. These gyres occur in the surface and thermocline regions of the ocean and are primarily controlled by the wind circulation, with modifications due to heating and cooling of the surface. The question now arises of why thermoclines are seen in the ocean. For example, why is the warm water not mixed over the whole depth of the ocean and why is the average ocean temperature about 31C. To explain the observed behavior thermohaline circulation, which is generated by small horizontal differences in density, due to temperature and salinity, between low and high latitude is considered. How does it work? Consider an ocean of uniform depth and bounded at the equator and at a polar latitude. We will assume it has initially a uniform temperature and is motionless (for the moment the effect of salinity on density are ignored). This hypothetical ocean is then subject to surface heating at low latitudes and surface cooling at high latitudes. In the lower latitudes the warming will spread downwards by diffusion, whereas in high latitudes the cooling will spread downwards by convection which is a much faster process than diffusion. The heavier colder water will induce a higher hydrostatic pressure at the ocean bottom than will occur at low latitudes. The horizontal pressure gradient at the ocean bottom is directed from the high latitudes to the lower latitudes, and will induce an equatorward abyssal flow of polar water. The flow can not move through the equatorial boundary of our hypothetical ocean and therefore will upwell into the upper layer of the tropical ocean, where it will warm by diffusion. The flow will then return polewards to the high latitudes where it will downwell into deepest layers of the ocean to complete the circuit. It is found that the downwelling occurs in narrow regions of the high latitudes whereas upwelling occurs over a very large area of the tropical ocean. This hypothetical ocean demonstrates the key role of the deep horizontal pressure gradient, caused by horizontal variations in density, for driving the flow. To explain the observed thermohaline circulation, this hypothetical ocean has to be modified to take into account the Coriolis force, which causes the deep abyssal currents to flow in narrow western boundary currents, the effect of salinity on the density (the haline component of the flow),
asymmetries in the buoyancy fluxes between the Northern and Southern Hemispheres and the complex bathemetry of the ocean basins. There follows a descriptive account of the thermohaline circulation. The deepest water masses (see Water Types and Water Masses) in the ocean have their origin in the polar seas. These seas experience strong cooling of the surface, particularly in the winter seasons. In the North Atlantic, there are connections through the Nordic seas to the Arctic Ocean, from which sea ice flows. Heat energy melts the ice in the North Atlantic and the melt water gives rise to further cooling. There are two effects on the density of the water: cooling increases the density whereas surface freshening, due to ice melt, decreases the density of the water. The former process usually dominates and hence denser waters are produced. These dense cold waters flow into the Atlantic through the East Greenland and West Greenland Currents and then into the Labrador Current. These cold waters mix and sink beneath the warm North Atlantic Current. In addition to surface polar currents there are also deep ocean currents (see Abyssal Currents). The cold saline water entering from the Nordic seas mixes as it sinks to the abyssal layers of the ocean and moves southward as a deep current along the western boundary of the Atlantic. This water mass is known as NADW (North Atlantic Deep Water) and it is the most prominent and voluminous of all the deep water masses in the global ocean. It flows into the Antarctic Circumpolar Current from where it flows into the Indian and Pacific Ocean. In addition to NADW, colder denser water, Antarctic Bottom Water (AABW) enters the Southern Ocean from the Antarctic shelf seas. It is not as voluminous as NADW but it flows northwards in the deepest layers into the Atlantic, where it can be distinguished as far north as 301N. These deep flows upwell into the thermocline and surface waters where they return to the North Atlantic. This global thermohaline circulation has been termed the global conveyor circulation to signify its role in transporting heat and fresh water (Figures 10 and 11) around the planet. How does this circulation explain the thermocline? The rate at which these cold deep abyssal waters are produced can be estimated and it is known for a steady state in the ocean that production has to be balanced by removal. A large-scale upwelling of the abyssal waters into the thermocline produces this removal. Our simple conceptual picture is of warm thermocline water mixing downwards, balanced by a steady upwelling of the cold abyssal layers. Without the upwelling, the warm waters would mix into the deepest layers of the ocean. (see Dispersion and Diffusion in the Deep Ocean).
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6 60˚N
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Figure 11 An estimate of the transfer of fresh water ( 109 kg s1) in the world ocean. In polar and equatorial regions precipitation and river run-off exceed evaporation, and hence there is an excess of fresh water, whereas in subtropical regions there is a water deficit. A horizontal transfer of fresh water is therefore required between regions of surplus to regions of deficit. FP and FA refer to the freshwater fluxes of the Pacific– Indian throughflow and of the Antarctic Circumpolar Current in the Drake Passage, respectively. (Reproduced from Wijffels SE, Schmitt R, Bryden H and Stigebrandt A (1992) Transport of freshwater by the oceans. Journal of Physical Oceanography 22: 155–162.)
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Figure 10 Poleward transfer of heat by (A) ocean and atmosphere together (TA þ TO), (B) atmosphere alone (TA), and (C) ocean alone (TO). The total heat transfer (A) is derived from satellite measurements at the top of the atmosphere, that of the atmosphere alone (B) is obtained from measurements of the atmosphere, and (C) is calculated as the difference between (A) and (B). (1 Petawatt (PW) ¼ 1015 W.) Data compiled from three sources. (Reproduced from Carrissimo BC, Oort AH and Von der Harr TH (1985) Estimating the meridional energy transports in the atmosphere and ocean. Journal of Physical Oceanography 15: 52–91.)
The Role of Fresh Water in Ocean Circulation The present discussion has shown that the winddriven circulation and the thermohaline circulation are major components of ocean circulation, which are ultimately driven by the surface wind stress and buoyancy fluxes. Buoyancy fluxes are the net effect of heat exchange and the freshwater exchange with the overlying atmosphere. It has been shown that heat exchange is a major process explaining existence of both the thermocline and the deep abyssal water but what is the role of the fresh water in ocean circulation? In the subtropics there is net removal of fresh water by evaporation. This increases the salinity of
the water which, in turn, increases the density of the thermocline waters. Normally this effect is opposed by heating, which lightens the water. However, in the Mediterranean and the Red Sea evaporation produces salient waters, which by virtue of their salinity and cooling in winter, sink to the deepest layers of the basins. At the Straits of Gibraltar (see Flows in Straits and Channels), this dense saline layer flows out beneath the incoming fresher and cooler Atlantic water. This Mediterranean water forms a distinct layer of high salinity water in the eastern Atlantic Ocean. Similar behavior occurs at Bab el Mandeb adjacent to the Gulf of Aden. The influence of fresh water is more substantial in the polar oceans. A given amount of fresh water will have a greater effect on density at low temperatures than at high temperatures, because the thermal expansion of sea water decreases with decreasing temperature. At higher latitudes there is a net addition of fresh water into the oceans, which arises from the excess of precipitation over evaporation and the melting of sea ice moving towards the equator from the polar regions. The addition of fresh water adds buoyancy to the surface layer while cooling removes buoyancy, therefore the fresh water will tend to reduce the effect of the cooling. In the Arctic Ocean (see Arctic Ocean Circulation) the surface layer is colder but less dense than the warmer layer at B100 m and therefore is in equilibrium. . This stable halocline in the Arctic Ocean reduces the vertical heat flux in to the deep ocean.
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OCEAN CIRCULATION
In the subpolar oceans, the addition of fresh water reduces the density of the surface layer and can reduce the prevalence of deep convection. This happened in the late 1960s when fresh water, probably from excessive ice in the Arctic Basin, melted in the subpolar gyre. The effect on the thermohaline circulation is unknown, but it is believed from modeling studies that the decrease in the production of deep waters reduced the thermohaline circulation of the ocean.
What are the Consequences of this Circulation on the Climate System? The effects of the ocean circulation on the climate can be understood in terms of the heat capacity of the ocean. The heat capacity of a column of sea water only 2.6 m deep is equivalent to that of a column of whole atmosphere and therefore the ocean heats and cools on a long timescale compared with the atmosphere. It is known that there is a poleward gradient of temperature, which is driven by the thermal radiation imbalance between the low and high latitudes. In response to this temperature gradient there is a flow of heat from the warmer to cooler latitudes. Both the atmosphere and ocean circulations transfer this heat from low to high latitudes by a variety of mechanisms. In the low latitudes of the atmosphere there is the Hadley cell which transfers low temperature air in the lowest levels via the trade winds towards the equator and transfers warmer air poleward in the upper troposphere (Figure 8). At higher latitudes anticyclones and cyclones and their accompanying upper air jet streams transfer heat polewards. In the ocean, the wind-driven Ekman currents transfer heat as surface waters move across latitude circles. This water is returned deeper in the ocean at a different temperature from that of the surface water. The ocean gyres carry heat towards higher latitudes since the poleward flows of the western part of the gyres are warmer than the equatorward flows in the eastern parts of the gyre. Finally, and not least, is the contribution of the thermohaline circulation, which transports warm surface and thermocline waters to the highest latitudes and returns cold water to lower latitudes. Figures 10 and 11 show the heat transport and fresh water transport in the ocean. A major difference between the atmosphere and ocean is the relative speed of their circulation. The atmosphere circulation is a fast system, responding on timescales of days to weeks. For example, weather systems in temperate latitudes grow and decay
on timescales of a few days. By contrast, the ocean tends to be slower in its response. The fastest part of the system are the surface Ekman layers which respond to changes in the surface wind circulation on a timescale of one or two days. Changes in wind circulation can cause planetary waves which will change sea level and surface temperature on monthly to seasonal timescales. In particular, the equatorial oceans respond to the surface wind stress on seasonal timescales, which allows a strong coupling between the ocean and atmosphere to take place. This gives rise to phenomena such as the El Nin˜o Southern Oscillation (see El Nin˜o Southern Oscillation (ENSO)). The subtropical gyres respond to changes in the wind circulation on decadal timescales, whereas the deep thermohaline circulation respond on millenial timescales. There is some evidence for rapid changes of local parts of the thermohaline circulation on timescales 50 years (see Abrupt Climate Change). Observations of the deep ocean are far fewer in number than at the ocean and land surface. The longest continuous data set is a deep station at 0
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Figure 12 Temperature changes (1C) in the subtropical North Atlantic (241N), 1957–1992. The measurements have been averaged across 241N between North Africa and Florida. (Reproduced from Parrilla G, Lavin A, Bryden H, Garcia M and Millard R (1994) Rising temperatures in the subtropical North Atlantic. Nature 369: 48–51.)
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OCEAN CIRCULATION
Bermuda that commenced operations in 1954. Observations from cruises in the earlier part of the century are of unknown quality and therefore it is difficult to know whether differences are due to the use of different instruments or to real changes in the ocean. It is only since the 1950s that such changes have been accurately measured. Figure 12 shows changes in the temperature for that period of time across the Atlantic. These changes are of the order of a few tenths of a degree over periods of 15 years. As the heat capacity of the oceans is very much larger than that of the atmosphere, these changes in temperature involve very significant changes in the heat content of the ocean. The World Ocean Circulation Experiment from 1990 to 1997 has provided measurements of ocean properties such as temperature, salinity, and chemical tracers as well as current measurements on a global scale. This set of high quality measurements will provide the baseline from which future changes in ocean circulation can be determined. Despite the brief record of deep ocean observations, sea surface temperature measurements of reasonable quality go back to the late nineteenth century. These measurements can be used to assess changes in the surface layers (see Heat Transport and Climate). Salinity measurements are fewer and not as reliable but, nevertheless, changes can be still detected (see Freshwater Transport and Climate). Salinity measurements in the Mediterranean over the last century have shown a warming of the Western Mediterranean Deep Water of 0.11C and increase of 0.05 in salinity. The reasons for this change are not known, but it has been speculated that the change in salinity may be attributed to a reduction in the freshwater flow due to the damming of the Nile and of rivers flowing into the Black Sea. An important recently identified question is the stability of the thermohaline circulation. The thermohaline circulation is driven by small density differences and therefore changes in the temperature and salinity arising from global warming may alter
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the thermohaline circulation. In particular, theoretical modeling of the ocean circulation has shown that the thermohaline circulation may be reduced or turned off completely when significant excess fresh water is added to the subpolar ocean. In the event of thermohaline circulation being significantly reduced or stopped, it may take many centuries before it returns to its present value (see Abrupt Climate Change, North Atlantic Oscillation (NAO)). In view of the current levels of uncertainty, it is necessary to continue to monitor the ocean circulation, as this will provide the key to the understanding of the present circulation and enhance our ability to predict future changes in circulation.
See also Abrupt Climate Change. Abyssal Currents. Antarctic Circumpolar Current. Arctic Ocean Circulation. CFCs in the Ocean. Current Systems in the Atlantic Ocean. Dispersion and Diffusion in the Deep Ocean. Drifters and Floats. El Nin˜o Southern Oscillation (ENSO). Elemental Distribution: Overview. Florida Current, Gulf Stream, and Labrador Current. Flows in Straits and Channels. General Circulation Models. Heat Transport and Climate. Inverse Models Moorings. North Atlantic Oscillation (NAO). Regional and Shelf Sea Models. Ships. Tritium–Helium Dating. Water Types and Water Masses. Waves on Beaches. Wind Driven Circulation.
Further Reading Gill AE (1982) Atmosphere–Ocean Dynamics. London: Academic Press. Siedler G, Church J, and Gould (2001) Ocean Circulation and Climate. London: Academic Press. Summerhayes CP and Thorpe SA (1996) Oceanography – An Illustrated Guide. London: Manson Publishing. Wells NC (1997) The Atmosphere and Ocean: A Physical Introduction, 2nd edn. Chichester: John Wiley.
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OCEAN CIRCULATION: MERIDIONAL OVERTURNING CIRCULATION J. R. Toggweiler, NOAA, Princeton, NJ, USA & 2009 Elsevier Ltd. All rights reserved.
Introduction The circulation of the ocean has been traditionally divided into two parts, a wind-driven circulation that dominates in the upper few hundred meters, and a density-driven circulation that dominates below. The latter was once called the ‘thermohaline circulation’, a designation that emphasized the roles of heating, cooling, freshening, and salinification in its production. Use of ‘thermohaline circulation’ has all but disappeared among professional oceanographers. The preferred designation is ‘meriodional overturning circulation’, hereafter MOC. This usage reflects the sense of the time-averaged flow, which generally consists of a poleward flow of relatively warm water that overlies an equatorward flow of colder water at depth. It also reflects a recognition by oceanographers that most of the work done to drive the circulation, whether near the surface or at depth, comes directly or indirectly from the wind. Like the thermohaline circulation of old, the MOC is an important factor in the Earth’s climate because it transports roughly 1015 W of heat poleward into high latitudes, about one-fourth of the total heat transport of the ocean/atmosphere circulation system. Radiocarbon measurements show that it turns over all the deep water in the ocean every 600 years or so. Its upwelling branch is important for the ocean’s biota as it brings nutrientrich deep water up to the surface. It may or may not be vulnerable to the warming and freshening of the Earth’s polar regions associated with global warming. The most distinguishing features of the MOC are observed in the sinking phase, when new deep-water masses are formed and sink into the interior or to the bottom. Large volumes of cold polar water can be observed spilling over sills, mixing violently with warmer ambient water, and otherwise descending to abyssal depths. The main features of the upwelling phase are less obvious. The biggest uncertainty is about where the upwelling occurs and how the upwelled deep water returns to the areas of deep-water formation.
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The Cooling Phase – Deep-water Formation The most vigorous overturning circulation in the ocean today is in the Atlantic Ocean where the upper part of the Atlantic’s MOC carries warm, upper ocean water through the Tropics and subtropics toward the north while the deep part carries cold dense polar water southward around the tip of Africa and into the Southern Ocean beyond. The Atlantic’s MOC converts roughly 15 106 m3 s 1 of upper ocean water into deep water. (Oceanographers designate a flow rate of 1 106 m3 s 1 to be 1 Sv. All the world’s rivers combined deliver c. 1 Sv of fresh water to the ocean.) The MOC in the Atlantic is often characterized as a ‘conveyor belt’ or more generally as a continuous current or ribbon of flow. It is shown following a path that extends through the Florida Straits and up the east coast of North America as part of the Gulf Stream. The ribbon then cuts to the east across the Atlantic and then northward closer to the coast of Europe. This is somewhat misleading because individual water parcels do not follow this kind of continuous path. The MOC is composed instead of multiple currents that transfer water and water properties along segments of the path. Individual water parcels loop around multiple times and may recirculate all the way around the wind-driven gyres while moving northward in stages. As the MOC moves northward through the tropical and subtropical North Atlantic, it spans a depth range from the surface down to B800 m and has a mean temperature of some 15–20 1C. During its transit through the Tropics and subtropics, the MOC becomes saltier due to the excess of evaporation over precipitation in this region. It also becomes warmer and saltier by mixing with the salty outflow from the Mediterranean Sea. By the time the MOC has crossed the 501 N parallel into the subpolar North Atlantic the water has cooled to an average temperature of 11–12 1C. Roughly half of the water carried northward by the MOC at this juncture moves into the Norwegian Sea between Iceland and Norway. Part of this flow extends into the Arctic Ocean. The final stages of this process make the salty North Atlantic water cold and dense enough to sink. Sinking is known to occur in three main places. The densest sinking water in the North Atlantic is
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formed in the Barents Sea north of Norway where salty water from the Norwegian Current is exposed to the atmosphere on the shallow ice-free Barents shelf. Roughly 2 Sv of water from the Norwegian Current crosses the shelf and sinks into the Arctic basin after being cooled down to 0 1C. This water eventually flows out of the Arctic along the coast of Greenland at a depth of 600–1000 m. The volume of dense water is increased by additional sinking and open ocean convection in the Greenland Sea north of Iceland. The dense water then spills into the North Atlantic over the sill between Greenland and Iceland in Denmark Strait (at about 600-m depth) and through the Faroe Bank Channel between the Faroe Islands and Scotland (at about 800-m depth). As 0 1C water from the Arctic and Greenland Seas passes over these sills, it mixes with 6 1C Atlantic water beyond the sills and descends the continental slopes down to a depth of 3000 m. The overflows merge and flow southward around the tip of Greenland into the Labrador Sea as part of a deep boundary current that follows the perimeter of the subpolar North Atlantic. A slightly warmer water mass is formed by open ocean convection within the Labrador Sea. The deep water formed in the Labrador Sea increases the volume flow of the boundary current that exits the subpolar North Atlantic beyond the eastern tip of Newfoundland.
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Newly formed water masses are easy to track by their distinct temperature and salinity signatures and high concentrations of oxygen. The southward flow of North Atlantic Deep Water (NADW) is a prime example. NADW is identified as a water mass with a narrow spread of temperatures and salinities between 2.0 and 3.5 1C and 34.9 and 35.0 psu, respectively. Figure 1 shows the distribution of salinity in the western Atlantic which tracks the southward movement of NADW. Newly formed NADW (S434.9) is easily distinguished from the relatively fresh intermediate water above (So34.6) and the Antarctic water below (So34.7). NADW exits the Atlantic south of Africa between 351 and 451 S and joins the eastward flow of the Antarctic Circumpolar Current (ACC). Traces of NADW reenter the Atlantic Ocean through Drake Passage (55–651 S) after flowing all the way around the globe. The other major site of deep-water formation is the coast of Antarctica. The surface waters around Antarctica, like those over most of the Arctic Ocean, are ice covered during winter and are too fresh to sink. Deep water below 500 m around Antarctica, on the other hand, is relatively warm (1.5 1C) and fairly salty (34.70–34.75 psu). This water mass, known as Circumpolar Deep Water (CDW), penetrates onto the relatively deep continental shelves around Antarctica where it is cooled to the freezing point.
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Figure 1 North–south section of salinity down the western Atlantic from Iceland to Drake Passage. The salinity distribution has been contoured every 0.1 psu between 34.0 and 35.0 psu to highlight NADW (34.9–35.0 psu).
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Brine rejection from the formation of new sea ice maintains fairly high salinities on the shelf despite the freshening effects of precipitation and the input of glacial meltwater. Very cold shelf water (–1 1C, 34.6–34.7 psu) is then observed descending the continental slope to the bottom in the Weddell Sea. Bottom water is also observed to form off the Adelie coast south of Australia (B1401 E). The volume of new deep water formed on the Antarctic shelves is not very large in relation to the volume of deep water formed in the North Atlantic, perhaps 3–4 Sv in total. It however entrains a large volume of old Circumpolar Deep Water, as it sinks to the abyss. Even with the entrainment, the deep water formed around Antarctica is denser than the NADW formed in the north. Antarctic Bottom Water (AABW) occupies the deepest parts of the ocean and is observed penetrating northward into the Atlantic, Indian, and Pacific Oceans through deep passages in the mid-ocean ridge system. Small quantities of deep water are also observed to form in semi-enclosed evaporative seas like the Mediterranean and Red Seas. These water masses are dense owing to their high salinities. They tend to form intermediate- depth water masses after exiting their formation areas in relation to the deep-water and bottom-water masses formed in the North Atlantic and around Antarctica.
The Warming Phase – Upwelling and the Return Flow The upwelling of deep water back to the ocean’s surface was thought at one time to be widely distributed over the ocean. This variety of upwelling was attributed to mixing processes that were hypothesized to be active throughout the interior. The mixing was thought to be slowly transfering heat downward, making the old deep water in the interior progressively less dense so that it could be displaced upward by the colder and saltier deep waters forming near the poles. Since the main areas of deep water formation are located at either end of the Atlantic, and since most of the ocean’s area is found in the Indian and Pacific Oceans, it stood to reason that the warming of the return flow should be widely distributed across the Indian and Pacific. Schematic diagrams often depict the closure of the old thermohaline circulation as a flow of warm upper ocean water that passes from the North Pacific through Indonesia, across the Indian Ocean, and then around the tip of Africa into the Atlantic. Observations made over the last 30 years point instead to turbulent mixing that is intense in some places but is not widespread. Attempts to directly
measure the turbulent mixing in the interior have shown that there may only be enough mixing to modify perhaps 10% or 20% of the deep water formed near the poles. There is no indication that any deep water is actually upwelling to the surface in the warm parts of the Indian and Pacific Oceans. Vigorous mixing is found near the bottom, where it is generated by tidal motions interacting with the bottom topography. It is also found in the vicinity of strong wind-driven currents. Thus, the energy that is available to warm the deep waters of the abyss comes mainly from the winds and tides. It now appears that most of the deep water sinking in the North Atlantic upwells back to the surface in the Southern Ocean. The upwelling occurs within the channel of open water that circles the globe around Antarctica. Figure 2 shows schematically how this is thought to work. The curved lines in the background are isolines of constant density (also known as isopycnals). Their downward plunge to the north away from Antarctica reflects the flow of the ACC out of the page in the center of the figure. Salty dense water from the North Atlantic is found at the base of the plunging isopycnals. Westerly winds above the ACC (also blowing out of the page) drive the ACC forward and also push the cold fresh water at the surface away from Antarctica to the north. Dense salty water from the North Atlantic is drawn upward in its place. A mixture of the salty dense deep water and the cold fresh surface water is then driven northward out of the channel by the westerly winds and is forced down into the thermocline on the north side of the ACC. In this way, the winds driving the ACC continually remove water with North Atlantic properties from the interior. Oceanographers are fairly certain that something like this is happening because the winds driving the ACC have become stronger over the last 40 years in response to global warming and the depletion of ozone over Antarctica. The subsurface water around Antarctica has become warmer, saltier, and lower in oxygen as upwelled water from the interior has displaced more of the cold fresh surface water, as shown in Figure 2. The thermocline water north of the ACC has also become cooler and fresher. Numerical experiments with ocean general circulation models show that much of the water forced down into the thermocline around the open channel eventually makes its way into the Atlantic Ocean where it is converted again into deep water in the northern North Atlantic. Thus, the winds over the ACC, in drawing up deep water from the ocean’s interior, have been shown to actively enhance the formation of deep water in the North Atlantic.
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OCEAN CIRCULATION: MERIDIONAL OVERTURNING CIRCULATION
Westerly winds
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Figure 2 Schematic cross section of the ACC, showing how the winds driving the current in an eastward direction out of the page also drive an overturning circulation. Stronger winds over the last 40 years have drawn more deep water to the surface south of the ACC, which has caused the subsurface waters around Antarctica to become warmer, saltier, and lower in oxygen, and have produced more downwelling to the north, which has made the thermocline waters north of the ACC cooler and fresher. Adapted from Aoki S, Bindoff NL, and Church JA (2005) Interdecadal water mass changes in the Southern Ocean between 30 1 E and 160 1 E. Geophysical Research Letters 32 (doi:10.1029/2005GL022220) and reproduced from Toggweiler RR and Russell J (2008) Ocean Circulation in a warming climate. Nature 451 (doi:10.1038/nature06590).
The deep water drawn up to the surface around Antarctica is colder than the surface water that is forced down into the thermocline north of the ACC. As this cold water comes into contact with the atmosphere and is carried northward, it takes up solar heat that otherwise would be available to warm the Southern Ocean and Antarctica. The MOC then carries this southern heat across the equator into the high latitudes of the North Atlantic where it is released to the atmosphere when new deep water is formed. The MOC thereby warms the North Atlantic at the expense of a colder Southern Ocean and colder Antarctica. Indeed, sea surface temperatures at 601 N in the North Atlantic are on average c. 6 1C warmer than sea surface temperatures at 601 S. If the warming phase of the MOC involved a downward mixing of heat in low and middle latitudes, as once thought, the MOC would carry tropical heat poleward into high latitudes. The warming phase of the Atlantic’s MOC now seems to take place in the south instead. This means that the heat transport by the MOC through the South Atlantic is directed equatorward and is opposed to the heat transport in the atmosphere. The net effect is a weakening of the global heat transport in the Southern Hemisphere and a strengthening of the heat transport in the Northern Hemisphere.
General Theory for the Meridional Overturning Circulation Figure 3 is a north–south section showing the distribution of potential density through the Atlantic Ocean. Most of the northward flow of the Atlantic’s MOC takes place between the 34.0 (B20 1C) and 36.0 (B8 1C) isopycnals, that is, within the main part of the thermocline, but some of the northward flow takes place among the denser isopycnals of the lower thermocline down to 36.6 g kg 1 (B1000 m). The southward flow of NADW is located, for the most part, below the 37.0 isopycnal. The isopycnals in the lower thermocline in Figure 3 are basically flat north of 401 S but rise up to the surface between 401 and 601 S. The rise of the isopycnals marks the eastward flow of the ACC, as shown previously in Figure 2. The relatively deep position of these isopycnals north of 401 S reflects the mechanical work done by the winds that drive the ACC. The convergent surface flow forced by the winds north of the ACC pushes the relatively light lower thermocline water down in relation to the cold, dense water that is drawn up by the winds south of the ACC. Numerical experiments carried out in ocean global climate models (GCMs) suggest that all the isopycnals of the lower thermocline would be squeezed up into the main thermocline if the
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Figure 3 Density structure of the Atlantic thermocline. Seawater density is referenced to a depth of 2000 m and has been zonally averaged across the Atlantic basin (units g kg 1). Contours have been chosen to highlight the lower part of the thermocline. The zonally averaged topography fails to capture the depths of the sills through the Greenland–Iceland–Scotland Ridge (651 N), which are at 600 and 800 m.
circumpolar channel were closed and the ACC eliminated. The relatively deep lower thermocline puts relatively warm water just to the south of the sills between Greenland, Iceland, and the Faroe Islands in the northern North Atlantic. This warm water is important because it sets up a sharp contrast with the cold water behind the sills, which ultimately drives the flow of dense water over the sills and into the deep Atlantic. In this way, the winds and the ACC in the south contribute as much, if not more, to the overflows in the north as the cooling that takes place at the surface. The depths of the isopycnals in the lower thermocline lead one to a general theory for the MOC in the Atlantic. Deep-water formation in the North Atlantic converts relatively low-density thermocline water into new deep water. It thereby removes mass from the thermocline and allows the isopycnals of the thermocline to be squeezed upward. The winds in the south have the opposite effect as they draw deep water up to the surface and pump it northward into the thermocline. They convert dense water from the deep ocean into low-density thermocline water and cause the thermocline to thicken downward. In this way, the thermocline thickness reflects a balance between the addition of mass via winds in the south and the loss of mass by deep-water formation in the north. The strength of the MOC should in this sense
be proportional to the thermocline thickness and the transformations in the north and south that convert light water to dense water and vice versa. A different kind of theory is needed for the overturning of the bottom water formed around Antarctica because the winds cannot play the same role. The upwelling branch of this circulation is also associated with the ACC as in Figure 2 but the upwelled water in this case flows poleward onto the adjacent Antarctic shelves where it is cooled and sinks back into the interior. The winds in this case help cool the upwelled water by exposing it to the atmosphere and by driving the fresh Antarctic surface waters away but the strength of the bottom water circulation would appear to be limited by the rate at which old bottom waters can be warmed by mixing with the overlying deep water.
Instability of the Meridional Overturning Circulation Cooling of the ocean in high latitudes makes polar surface waters denser in relation to warmer waters at lower latitudes. Thus cooling contributes to a stronger MOC. The salinity section through the Atlantic Ocean in Figure 1 gives one a superficial impression that salinification also makes a positive contribution to the MOC. This is actually not true.
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The cycling of fresh water between the ocean and atmosphere (the hydrological cycle) results in a net addition of fresh water to the polar oceans which reduces the density of polar surface waters. Thus, the haline part of the old thermohaline circulation is nearly always in opposition to the thermal forcing. The Earth’s hydrological cycle is expected to become more vigorous in the future with global warming. This may weaken the MOC in a way which could be fairly abrupt and unpredictable. NADW is salty because the upper part of the MOC flows through zones of intense evaporation in the Tropics and subtropics. If the rate at which new deep water is forming is relatively high, as it seems to be at the present time, the sinking water removes much of the fresh water added in high latitudes and carries it into the interior. The added fresh water in this case dilutes the salty water being carried into the deep-water formation areas but does not erase the effect of evaporation in low latitudes. Thus, NADW remains fairly salty and is able to export fresh water from the Atlantic basin. If the rate of deep-water formation is relatively low or the hydrological cycle is fairly strong, the fresh water added in high latitudes can create a low-salinity lid that can reduce or annihilate the MOC. There seems to be a critical input of freshwater input for maintaining the MOC. If the freshwater input is close to this threshold, the overturning becomes unstable and may become prone to wild swings over time. Coupled (ocean þ atmosphere) climate models from the 1990s projected that a fourfold increase in atmospheric CO2 would increase the hydrological cycle sufficiently that the MOC might collapse. More recent coupled models are predicting that the warming will, in addition, lead to stronger mid-latitude westerly winds that are shifted poleward with respect to their preindustrial position. This change in the westerlies puts stronger Southern Hemisphere westerlies directly over the ACC, which should make the MOC stronger. Thus, the MOC could become stronger or weaker depending on whether the winds or the hydrological cycle dominate in the future.
See also Antarctic Circumpolar Current. Florida Current, Gulf Stream, and Labrador Current. Internal Tidal Mixing. Ocean Circulation. Ocean Subduction. Upwelling Ecosystems. Water Types and Water Masses. Wind Driven Circulation.
Further Reading
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30 1 E and 160 1 E. Geophysical Research Letters 32: (doi:10.1029/2005GL022220). Broecker WS (1991) The great ocean conveyor. Oceanography 4: 79--89. Gnanadesikan A (1999) A simple predictive model for the structure of the oceanic pycnocline. Science 283: 2077--2079. Kuhlbrodt T, Griesel A, Montoya M, Levermann A, Hofmann M, and Rahmstorf S (2007) On the driving processes of the Atlantic meridional overturning circulation. Reviews of Geophysics 45: RG2001 (doi:10.1029/2004RG000166). Manabe S and Stouffer R (1993) Century-scale effects of increased atmospheric CO2 on the ocean–atmosphere system. Nature 364: 215--218. McCartney MS and Talley LD (1984) Warm-to-cold water conversion in the northern North Atlantic Ocean. Journal of Physical Oceanography 14: 922--935. Munk W and Wunsch C (1998) Abyssal recipes II: Energetics of tidal and wind mixing. Deep-Sea Research I 45: 1977--2010. Orsi A, Johnson G, and Bullister J (1999) Circulation, mixing, and production of Antarctic bottom water. Progress in Oceanography 43: 55--109. Rahmstorf S (1995) Bifurcations of the Atlantic thermohaline circulation in response to changes in the hydrological cycle. Nature 378: 145--149. Rudels B, Jones EP, Anderson LG, and Kattner G (1994) On the intermediate depth waters of the Arctic Ocean. In: Johannessen OM, Muench RD, and Overland JE (eds.) Geophysical Monograph 85: The Polar Oceans and Their Role in Shaping the Global Environment, pp. 33--46. Washington, DC: American Geophysical Union. Russell JL, Dixon KW, Gnanadesikan A, Stouffer RJ, and Toggweiler JR (2006) The Southern Hemisphere westerlies in a warming world: Propping open the door to the deep ocean. Journal of Climate 19(24): 6382--6390. Schmitz WJ and McCartney MS (1993) On the North Atlantic Circulation. Reviews of Geophysics 31: 29--49. Toggweiler JR and Samuels B (1995) Effect of Drake Passage on the global thermohaline circulation. Deep-Sea Research I 42: 477--500. Toggweiler JR and Samuels B (1998) On the ocean’s largescale circulation near the limit of no vertical mixing. Journal of Physical Oceanography 28: 1832--1852. Toggweiler RR and Russell J (2008) Ocean Circulation in a warming climate. Nature 451: (doi:10.1038/nature 06590). Tziperman E (2000) Proximity of the present-day thermohaline circulation to an instability threshold. Journal of Physical Oceanography 30: 90--104. Wunsch C and Ferrari R (2004) Vertical mixing, energy, and the general circulation of the oceans. Annual Reviews of Fluid Mechanics 36: 281--314.
Aoki S, Bindoff NL, and Church JA (2005) Interdecadal water mass changes in the Southern Ocean between
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OCEAN GYRE ECOSYSTEMS M. P. Seki and J. J. Polovina, National Marine Fisheries Service, Honolulu, HI, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1959–1965, & 2001, Elsevier Ltd.
The Generalized Open-ocean Food Web Food webs, simply put, describe all of the trophic relationships and energy flow between and among the component species of a community or ecosystem. A food chain depicts a single pathway up the food web. The first trophic level of a simple food chain in the open ocean begins with the phytoplankton, the autotrophic primary producers, which build organic materials from inorganic elements. Herbivorous zooplankton that feed directly on the phytoplankton are the primary consumers and make up the second trophic level. Subsequent trophic levels are formed by the carnivorous species of zooplankton that feed on herbivorous species and by the carnivores that feed on smaller carnivores, and so on up to the highest trophic level occupied by those adult animals that have no predators of their own other than humans; these top-level predators may include sharks, fish, squid, and mammals. In ocean gyres, the dominant phytoplankton, especially in oligotrophic waters, are composed of very small forms, marine protozoans such as zooflagellates and ciliates become important intermediary links, and the food chain is lengthened. There are thus typically about five or six trophic levels in these ecosystems. In contrast, large diatoms dominate in nutrient-rich upwelling regions, resulting in shorter food chains of three or four trophic levels since large zooplankton or fish can feed directly on the larger primary producers. Production of larger flagellates/diatoms in specialized habitats of the open ocean may lead to shortened energy paths. As there are seldom simple linear food chains in the sea, a food web with multiple and shifting interactions between the organisms involved portrays more accurately the trophic dynamics of a given ecosystem. Examples of food webs are presented for the North Pacific Subarctic and Subtropical gyres in Figures 1 and 2, respectively. The place a particular species occupies in the ecosystem food web is not necessarily constant, since feeding requirements change as organisms grow. Some species change diets
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(and trophic levels) as they grow or as the relative abundance and availability of different food items change. In some species, cannibalism of their own young may be important. Before proceeding, we need to recognize that in ocean gyre ecosystems, very large percentages of organic matter are cycled through microbes before entering the linear arrangement of the classic food web. The role of viruses, bacteria, heterotrophic nanoflagellates, nano- and microplanktonic protozoans, and the microbial loop in the open ocean ecosystem is discussed in detail elsewhere in the Encyclopedia and we only note here that if several trophic steps are involved in a microbial food web, there must be a significant loss of carbon at each transfer and therefore little transfer of carbon from microbial to classic planktonic food webs.
The Ocean Gyre Ecosystem In each of the major ocean basins, surface winds drive currents that form the large anticyclonic subtropical gyres as well as the smaller, cyclonic subarctic gyres. The details of both vertical structure and trophic links in the ocean gyre food web differ regionally and seasonally; opportunism is generally the rule of prey selection and is driven by regional and seasonal differences in vertical physical structure of the upper water column. The surface waters in the warm subtropical gyres away from seasonal meridional frontal systems tend to be highly stratified, with a permanent thermocline that limits vertical enrichment of the euphotic zone throughout the year. In this nutrient depleted environment, recycled nitrogen (primarily in the form of ammonia and urea) excreted by the zooplankton or released by the microbial community provides the primary nitrogen for a continued low level of phytoplankton ‘regenerated’ production that remains in approximate balance by zooplankton grazing. These oligotrophic waters are characteristically recognized as some of the least productive waters in the world’s oceans. Alternatively, ‘new’ primary production based on the input of new nitrogen (primarily nitrate) into the euphotic zone occurs at much lower levels through oceanographic physical forcing, atmospheric input, or nitrogen fixation. As in most other ocean systems, a deep chlorophyll maximum (DCM) is present, but this will vary spatially and temporally with consequent
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Subarctic Food Web Trophic Level
Northern fur seals
Toothed whales porpoises
6
Daggertooth
Murres shearwaters Baleen whales
Elephant seal
Salmon shark
Large nektonic squid
Pomfret
5 Chum salmon
Chinook salmon steelhead trout
Sockeye salmon
Pink salmon
Coho salmon
Saury 4
Micronektonic squid
Ctenophores Sergestid shrimp
Myctophids bathlyagids
Heteropods salps
Hyperiids 3 Pteropods
Euphausiids
Copepods
Larvaceans
2
1
Microzooplankton
Phytoplankton
Figure 1 Food web of the Subarctic North Pacific oceanic ecosystem; arrows point in the direction of energy flow. (Adapted from Brodeur et al. (1999).)
changes in the depth of the thermocline, nutrient flux, and the maximum primary production rate. In contrast, subarctic oceans benefit from alternating periods of summer stratification and winter mixing, so that there is a seasonal injection of nutrients into surface waters. In the North Atlantic the trophic framework is set up by the classic temperate bloom cycle; i.e., a ‘spring bloom’ followed by a ‘crash’, a ‘stable state summer equilibrium’, a smaller ‘fall bloom’, and a ‘winter decline’. Winter storms create deeper well-mixed conditions including the presence of high nitrate levels. With the onset of spring comes an increase in solar irradiation and water column stability. Phytoplankton (normally diatoms) ‘blooms’ with nutrient injection. Slightly lagging behind the phytoplankton bloom is a bloom of zooplankton that feeds on the increased phytoplankton. As the increased phytoplankton population exhausts the nitrate supply and zooplankton grazing overshoots phytoplankton growth, a ‘crash’ occurs, returning nitrate, phytoplankton, and zooplankton to low, stable levels. During the summer months, water is well stratified with balanced processes (i.e., phytoplankton growth (limited by lower levels of available nitrate) is balanced by zooplankton grazing). During
fall, early episodic storms induce some mixing and hence phytoplankton blooms followed by periods of stability. The onset of increased winter storm activity mixes nutrients in surface waters; however, available light energy decreases, limiting phytoplankton production. Thus the low phytoplankton production periods are light limited during the winter and nutrient-limited during the summer. In the subarctic North Pacific there is virtually no bloom, diatoms populations are low, and high abundances of microzooplankton graze the small phytoplankton, keeping the abundance down. The subarctic North Pacific waters are thus high in nutrients (nitrogen) but low in chlorophyll (HNLC), similar to those ecosystems found near the equator and in the Antarctic. Modest winter mixing in the subarctic Pacific results in a more favorable light regime and a continued modest level of biological production throughout the year and reduced biomass accumulation that is less strongly pulsed than that occurring in the Atlantic. The constant planktonic biomass in the Pacific favors pelagic fish production in contrast to the Atlantic, where much of the plankton biomass is lost to sinking without being grazed, thereby supporting more benthic resources.
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OCEAN GYRE ECOSYSTEMS
Subtropical Food Web
6
Yellowfin & Skipjack tuna
Terns & Shearwaters
5
Bigeye tuna
Mahimahi
Trophic level
Baleen whales
Toothed marine mammals
Sharks
Billfishes
Sea turtles
Albatrosses
Pomfrets, nomeids
Other squid
Flying squid
Flyingfish, carangids
Myctophids
4
3
Heteropods pteropods
Decapod crustaceans
Micronektonic squid
Euphausiids
Copepods
2
Gelatinous zooplankton
Microzooplankton
1
Phytoplankton
Figure 2 Food web of the Subtropical North Pacific oceanic ecosystem; arrows point in the direction of energy flow.
The Intermediate Trophic Levels The animals occupying the intermediate trophic levels that link the primary producers and consumers with the top of the food web include all of the zooplankton and micronekton (smaller organisms 10– 100 cm in length) large enough to swim in inertial conditions. Two groups of animals in particular play a key role in the ocean gyre food web: those that compose the vertically migrating deep scattering layer (DSL) and the small pelagic ‘forage fishes’. As night approaches, myriads of animals make an ascent from various depths to grazing or hunting grounds near the surface. Before daybreak, the DSL organisms descend back to their deeper, daytime residence. These migrations range 200–700 m in vertical extent and are confined to the epipelagic and mesopelagic zones. With the diel DSL migration, concentrations of food near the surface are much greater (a hundredfold increase in population numbers) at night than during day, establishing a forage base for larger, aggressive, nocturnally active carnivores; e.g., dolphins and pelagic squid. In the tropics, large filter-feeding organisms such as whale sharks, baleen whales, and
megamouth sharks probably also utilize these highdensity layers of plankton and micronekton. Diel vertical movements of bigeye tuna (Thunnus obesus) exhibit high correlation with the migration of the DSL and their associated prey species. The DSL micronekton, which includes fishes (e.g., myctophids, gonostomids), squid, and crustaceans (especially euphausiids and sergestid shrimps) are likely the primary zooplanktivores in the open ocean. A large component of the DSL faunal composition is the gelatinous zooplankton (combjellies, pelagic tunicates, medusae, ctenophores, siphonophores, etc.). Despite their ubiquitous and abundant nature, the significance of these zooplankton in the pelagic food web has been and still is somewhat equivocal. These gelatinous animals had been thought to represent an ecological dead end. Because of their low food value, energy taken up through their phytoplankton grazing was not passed up to higher trophic levels. From the growing list of fish, sea turtle, and seabird species identified as habitually feeding on gelatinous prey, there is strong evidence that this resource plays more of a key role in the gyre food web than previously believed.
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In the open ocean, small pelagic ‘forage fishes’ are the keystone prey species among apex predators. By definition, forage fishes are ‘y abundant, schooling fishes (and squid) preyed upon by many species of seabirds, marine mammals, and other fish species’ according to A. M. Springer and S. G. Speckman. These animals are generally short-lived, highly fecund, and heavily preyed on as juveniles as well as adults. In the subtropical gyre, loose aggregations of flying squid, pomfret, saury, or mackerel, and in temperate waters tighter schools of anchovies, sardines, or herring, are representative of this guild. From a trophic standpoint, the importance of pelagic ommastrephid squids in the gyre ecosystems cannot be understated. These forage animals have rapid population turnover rates and can dominate the epipelagic nektonic biomass, as is the case with the red flying squid Ommastrephes bartramii at the summer Subarctic Boundary. Throughout the gyre food webs, ommastrephids often dominate the diets of toothed cetaceans, pinnipeds, fishes, and seabirds, as well as of tropical seabirds and marine mammals.
The Higher Trophic Levels The faunal compositions of epipelagic nekton communities tend to be more diverse and more complex in the subtropics than those found in the temperate subarctic. At the top levels of the epipelagic (0–400 m) food web (and particularly within the 0–100 m euphotic zone) are the fast-swimming predators such as sharks, billfishes, and large tunas as well as some of the marine mammals and surfacefeeding seabirds that feed at the fourth and fifth trophic levels. Generally, most of these breed in warmer tropical fall/winter waters, some seabirds notwithstanding, but make major excursions to the northern subtropics–subarctic during spring/summer to feed; all are opportunistic predators within their respective niches. Large pelagic fishes are the most abundant and diversified of the apex predators but are nevertheless still highly dispersed, with typical population densities of only one fish per square kilometer integrated over the upper mixed layer. Blue sharks (Prionace glauca) are by far the most abundant shark species in the ocean gyres and in subtropical waters; frequent encounters with fishing gear make them seem ubiquitous. Dominant prey items for blue shark include the abundant ommastrephid squids and forage fishes. Billfishes, including the marlins (Makaira spp. and Tetrapturus spp.) and swordfish (Xiphias gladius), are capable of extensive vertical excursions of several hundred meters but concentrate feeding at the
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surface on smaller fishes and, in the case of swordfish, exhibit a particular predilection for pelagic squid. Tunas (Scombridae) are diurnal predators and feed from the surface down to about 400 m. With the exception of bigeye tuna, tunas feed almost solely on small surface dwelling fauna – i.e., they rarely feed on the vertically migrating DSL. Juvenile and postlarval crustaceans, cephalopods, and oceanic- (Bramidae, Nomeidae, etc.) and shore-fishes (Mullidae, Chaetodontidae, etc.) dominate the tuna diet. Insular-based prey are likely artifacts of the inherent bias introduced from land-based fishing operations that underlie our present perception of oceanic predatory fish diets; these diminish in importance to the food web with distance from land. Among marine mammals, baleen whales are filter feeders of large zooplankton and hence feed at the third trophic level. Toothed whales and delphinids all feed at the fifth or sixth trophic level on pelagic cephalopods and, depending on the respective species’ distributional range, on epipelagic and mesopelagic fish. In the North Pacific, only two species of pinnipeds, the northern fur seal and northern elephant seal use the open ocean regions of the gyres to any extent. For these mammals, stomach contents suggest that oceanic squid are the key food web link. Most seabirds foraging in the high seas of the ocean gyres are Procellariiformes, namely the albatrosses, shearwaters, petrels, and fulmars. These seabirds are opportunistic surface feeders on small ‘forage fishes’. In the oligotrophic subtropics close to nesting colonies, a feeding guild of shearwaters and terns represents numerically the most abundant among breeding seabird species, relying on predatory fish such as tuna to drive schools of small pelagics to the ocean surface, thereby facilitating feeding.
Enhanced Feeding Habitats The pelagic zone of the open ocean gyre is often perceived as the most monotonous living space of our planet, with few visual cues to maintain spatial orientation, and spatial heterogeneity effectively restricted to vertical gradients of light, temperature, and abundance of organisms. However, regions of horizontal oceanographic variability abound throughout the gyre ecosystem in the form of largescale frontal systems and mesoscale dynamic features where productivity is enhanced and trophic transfer facilitated. A discussion of the gyre food web would not be complete without highlighting these dynamic features where so much of the ecosystem energy is mobilized.
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OCEAN GYRE ECOSYSTEMS
Large-scale fronts and associated frontal zones in the open ocean have profound effects on the distribution and movement patterns of pelagic animals. Generally, the distribution of pelagic animals tends to be coupled with preferred thermal habitats, a response to forage accumulation (enhanced forage opportunity), migration cues, and/or energetic gains from riding currents. On one hand, these basin-scale features form the boundaries that divide some of the large, core pelagic biogeographic provinces; on the other, the fronts are recognized as regions of convergence where life-forms on all trophic levels are concentrated and so support feeding and spawning aggregations. In the North Pacific, these thermohaline fronts mark the bounds for many of the Transition Zone keystone species, including the Pacific pomfret (Brama japonica), the red flying squid, and blue shark, which undergo extensive seasonal migrations northward during summer to feed along the Subarctic Boundary convergence and southward during the winter and spring to spawn in the subtropics. The aggregation of all trophic levels at the Subarctic Boundary results in a complex but trophically rich community where considerable energy is transferred. During the winter–spring, shoaling of the thermocline at the North Pacific Subtropical Front is believed to elevate the nutricline into the euphotic zone, resulting in enhanced local productivity, particularly at the depth of the DCM, as well as physically concentrating loose aggregations of prey closer to the ocean surface for predation. A basinwide surface chlorophyll front, the Transition Zone Chlorophyll Front, is also manifested in the surface waters of the North Pacific. This biological front lies at the boundary between the low-chlorophyll subtropical and high-chlorophyll subarctic gyres. Recently, neuston-feeding loggerhead sea turtles, Caretta caretta, have been found to exhibit strong affinities for this specialized habitat. Mesoscale variability on spatial scales of 10– 100 km – the ‘weather’ of the ocean – includes highgradient dynamic features such as frontal meanders, eddies, and jets. These features often give rise to localized regions of higher productivity, leading to aggregation and development of a forage base while physical gradients provide cues for predators to locate prey or more directly aggregate or concentrate food items. In stratified, oligotrophic waters of the large ocean gyres, recycling of nutrients between the grazers and the phytoplankton typically maintains primary production at uniformly low levels. Transient episodes of upwelled nutrient-rich water from mesoscale events, particularly strong cyclonic eddies and meanders, have been shown to induce ‘new’ production, thus providing a mechanism to shorten the trophic
pathway and facilitate energy transfer. The injection of nutrients at these features increases the biomass of phytoplankton through the contribution of larger eukaryotic phytoplankton, notably diatoms and dinoflagellates, over the recycled-nutrient-based picophytoplankton (e.g., photosynthesizing cyanobacteria species) that normally typify the system. Seamounts, also known as submarine rises or table mounts, can have a strong influence on adjacent openocean food webs in a variety of ways, particularly those that rise within the upper few hundred meters of the surface. Waters overlying seamounts are often characterized by high standing stocks of plankton and, at some locations, concentrate and transfer energy not only among the pelagic community but also to the demersal resident ecosystems. Several phenomena contribute to the maintenance and enhancement of the unique seamount communities. For one, some seamount ecosystems possess midwater micronekton communities that include both a unique resident assemblage as well as a DSL component advected from adjacent waters. At shallower seamounts, the members of the DSL community rise in the water column at night, are advected over the seamount, and are subsequently trapped on the summit as they descend at dawn, creating an enhanced feeding regime during the morning hours. The resident faunal assemblages are analogous to the land-associated mesopelagic boundary communities identified on larger island and shelf slopes. Seamounts are also hypothesized to benefit from the development of Taylor columns, or semistationary eddies located above the seamount, that would help retain planktonic populations and enhance productivity.
Food Web Dynamics Food webs of ocean gyres are altered by a range of factors, including fishing and variations in physical forcing resulting from climate fluctuations. The relationship between climate variation and ecosystem dynamics has been a focus of considerable research. Changes in abundance of organisms at many trophic levels have been documented to vary coherently with various atmospheric indices, including the Aleutian Low in the Pacific and the North Atlantic Oscillation in the Atlantic. For example, in the North Pacific subarctic gyre, changes in the fishery catches of all species of salmon vary on timescales of decades, coherent with the intensity of the Aleutian Low Pressure System. While the link between the atmosphere and salmon productivity has not been resolved, a number of hypotheses suggest that changes in salmon carrying capacity result from changes in
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OCEAN GYRE ECOSYSTEMS
physical forcing that impact productivity at the bottom of the food web and propagate up the food web to salmon. A change in the energy flow through the food web from the bottom up in response to climate variation is typical of many relationships described between climate and food web dynamics in ocean gyres. Fisheries in ocean gyres typically impact food webs through removals at top or mid-trophic levels. For example, longline or troll fisheries target apex predators including tunas, swordfish, marlins, and salmon, although some mid-trophic level species including squid can also be the target species. By-catch and incidental catches in ocean gyre fisheries also include apex species including sharks and protected species such as seabirds, marine mammals, and sea turtles. Data on the responses of oceanic gyre food webs to fishing are generally limited to harvested species, so the food web impacts at lower trophic levels are not documented. However, a dynamic ecosystem model has been used to investigate possible impacts and has found no evidence that the removal of any single high trophic level species significantly altered the food web. The lack of a keystone species appears to be due to a high degree of diet overlap among the high trophic level species. Fisheries in oceanic gyres alter the food web by reducing biomass at the top of the food web. When this reduction becomes substantial, it may result in some increase in biomass at mid-trophic levels.
See also Ecosystem Effects of Fishing. Fish Feeding and Foraging. Fish Predation and Mortality. Fish Migration, Vertical. Large Marine Ecosystems. Marine Mammal Trophic Levels and Interactions. Mesopelagic Fishes. Ocean Gyre Ecosystems. Pelagic Fishes. Seabird Foraging Ecology.
Further Reading Boehlert GW (1986) Productivity and population maintenance of seamount resources and future research
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directions. In: Uchida RN, Hayasi S, and Boehlert GW (eds) Environment and Resources of Seamounts in the North Pacific, pp. 95–101 Washington, DC: US Department of Commerce. NOAA Technical Report NMFS 43. Brodeur R, McKinnell S, Nagasawa K, et al. (1999) Epipelagic nekton of the North Pacific Subarctic and Transition Zones. Progress in Oceanography 43: 365– 397. Francis RC, Hare CR, Hollowed AB, and Wooster WS (1998) Effects of interdecadal climate variability on the oceanic ecosystems of the NE Pacific. Fisheries Oceanography 7: 1--21. Josse E, Bach P, and Dagorn L (1988) Simultaneous observations of tuna movements and their prey by sonic tracking and acoustic surveys. Hydrobiologia 371/372: 61--69. Kajimura H and Loughlin TR (1988) Marine mammals in the oceanic food web of the eastern subarctic Pacific. In: Nemoto T and Pearcy WG (eds.) The Biology of the Subarctic Pacific. Bulletins of the Ocean Research Institute, 187--223. Kitchell JF, Boggs CH, He X, and Walters CJ (1999) Keystone predators in the central Pacific. Ecosystem Approaches for Fisheries Management, pp. 665--683. Fairbanks, AK: Alaska Sea Grant College. Longhurst AR and Pauly D (1987) Ecology of Tropical Oceans. San Diego: Academic Press. Mann KH and Lazier JRN (1991) Dynamics of Marine Ecosystems. Biological–Physical Interactions in the Oceans. Cambridge, MA: Blackwell Scientific. Olson DB, Hitchcock GL, Mariano AJ, et al. (1984) Life on the edge: marine life and fronts. Oceanography 7: 52–60. Polovina JJ (1994) The case of the missing lobsters. Natural History 103(2): 50--59. Polovina JJ, Howell E, Kobayashi DR and Seki MP (2001) The Transition Zone Chlorophyll Front, a dynamic, global feature defining migration and forage habitat for marine resources. Progress in Oceanography in press. Springer AM and Speckman SG (1997) A forage fish is what? Summary of the symposium. In: Proceedings. Forage Fishes in Marine Ecosystems, pp. 773--805. Fairbanks, AK: Alaska Sea Grant College. Steele JH (1974) The Structure of Marine Ecosystems. Cambridge, MA: Harvard University Press. Valiela I (1995) Marine Ecological Processes. New York: Springer-Verlag.
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OCEAN MARGIN SEDIMENTS S. L. Goodbred Jr, State University of New York, Stony Brook, NY, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 4, pp 1965–1973, & 2001, Elsevier Ltd.
Introduction Ocean margin sediments are largely detrital deposits of terrestrial origin that extend from the shoreline to the foot of the continental rise. Indeed, about 80% of the world’s sediment is stored within margin systems, which cover about 14% of the Earth’s surface (Table 1; Figure 1). Margin deposits typically consist of sand, silt, and clay-sized particles, the characteristics of which reflect the geology and climate of the adjacent continent. Some of the signals relevant to terrestrial conditions include sediment size, mineralogy, geochemistry, and isotopic signature. Upon entering the marine realm, however, sediments take on new characteristics indicative of coastal ocean processes that include waves, tides, currents, sea level, and biological productivity. Given that these numerous terrestrial and marine processes impart a signature to the sediments, ocean margins preserve an important record of Earth history, providing insights into past atmospheric, terrestrial, and marine conditions. Beyond their significance as environmental recorders ocean margin sediments support major petroleum and mineral resources. Currently, over 50% of the world’s oil is recovered along ocean margins, and much of the remaining fraction is held within ancient margin deposits. In the 1950s, early investigations of ocean margins focused on tectonic structure and how overlying sedimentary sequences developed on timescales of Table 1 Area of Earth’s major physiographic provinces and the volume of stored sediments. Bracketed are ocean margin components with relative contributions shown
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105–107 years. Originally aimed at understanding plate tectonics and the nature of ocean–continent boundaries, these large-scale studies continued through the 1960s with specific interest in petroleum resources. Such efforts culminated in the publication of large scientific volumes such as Burk and Drake’s Geology of Continental Margins (1974). At the time of these summary publications, new models of sedimentary margin systems were already being developed, has also been notably the approach of seismic stratigraphy established at the Exxon Production Research Company. Growing out of this approach was the more general model of sequence stratigraphy, which helped establish that margin strata could be grouped into discrete packages reflecting cycles of sea-level rise and fall. These concepts represent a general approach that has been applied to stratigraphic development and margin evolution in most of the world’s ancient and modern sedimentary systems. Sequence stratigraphic data has also been mated with lithologic, magnetic, and biostratigraphic records, allowing researchers to establish the history of sea-level change since the Triassic era (260 million years ago). This historic record significantly advanced out understanding of ocean margin sedimentary records, because sea level is a major control on the distribution and accumulation of margin deposits, as well as an indicator of global climate. At a much shorter time scale, a great deal of research in recent decades has focused on sediment dynamics along modern margin systems. Aimed at understanding the fate of sediments entering the marine realm, some of the issues driving marginsediment research included coastal hazards, land loss and development, storm impacts, contaminant fate, and military interests. In part, the field has advanced with the development of new technology such as marine radioisotope geochronology, sonar seafloor mapping, and instruments for remote wave, current, and sediment concentration measures. Recently, research programs have sought to integrate long and short geologic timescales to understand how daily, seasonal, and annual sedimentary beds are ultimately preserved to form thick millennial and longer sedimentary sequences (e.g., US and European STRATAFORM programs). Other newer initiatives are seeking to take integration a step further by recognizing that ocean margin sediments lie among a continuum of terrestrial and marine influences. Therefore, to better understand this critical accumulation zone, the
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Figure 1 Distribution of major ocean margin deposits, showing extent of continental shelves (black) and rises (hatched). (Modified from Emery (1980).)
terrestrial processes that are responsible for sediment input must be considered in conjunction with the marine processes that are responsible for their redistribution (e.g., US and Japanese MARGINS programs). Such efforts to integrate over various spatial and temporal scales will likely govern the foreseeable future of margin sedimentary research.
Margin Structure All margins represent past or present tectonic plate boundaries where crusts of varying age, density, and composition meet. In general these boundaries also comprise large-scale basins that trap sediments shed from the adjacent land surface. Most ocean margins specifically denote the boundary between oceanic and continental crusts, where different densities between these adjacent blocks give rise to a steep gradient at the margin. This tectonic structure controls the physiography of the margin as well, which comprises shelf, slope, and rise settings (Figure 2). The landward part of the margin, the shelf, overlies the buoyant continental-crust block and is thus the shallowest part of the margin, extending along a low gradient (0.11) from the coast to the shelf break at 100–200 m water depth. Because of the low gradient and shallow depths, shelf environment are sensitive to sea-level change and isostatic movements (crustal buoyancy) causing the shelf to be periodically flooded and exposed. The seaward edge of the shelf, called the break, is identified by an increase in seafloor gradient (3–61) and represents a transition between the shelf and continental slope. Roughly overlying the continental and oceanic crusts, sediment accumulation on the slope is partly controlled by on the shelf, because any slope sediments must first bypass this depositional region. Beyond the slope extends the rise, a
low gradient (0.3–1.41) sedimentary apron that begins in 1500–3500 m water depth. Although the rise may extend hundreds of kilometers into the ocean basin, its sediment source largely derives from sediment originating on the slope and moves as turbidity currents. In addition to the major shelf, slope, and rise features, other regionally significant marginal features may include deltas and canyons, each playing a role in controlling or modifying the dispersal and deposition of sediments. Tectonically, margins may be broadly divided into divergent (passive) and convergent (active) systems, each with a characteristic structure and physiography that are important to sediment accumulation and transport dynamics. Divergent margins are largely stable crustal boundaries, although the structure and movement of deep basement rocks remain a significant influence on sediment deposition and stratigraphy. Divergent margins are also largely constructive settings in which sediments shed from the continent form thick sedimentary sequences and a wide, low gradient shelf. This loading of sediment onto the margin also drives a slow downwarping of the crust, creating more accommodation space for the accumulation of thick sedimentary sequences. The resulting margin physiography is characterized by a broad shelf, slope, and rise. Margins along the Atlantic Ocean are among the best-studied examples of divergent systems. In contrast, convergent margins are characterized by the subduction of oceanic crust beneath the adjacent continental block. Active mountain building in convergent settings frequently drives uplift of the margin, limiting the development of thick sedimentary sequences on the shelf. Thus, convergent margins tend to have narrow shelves that bypass most sediment to a steep and well-developed slope. The rise is generally not a significant depocenter because
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OCEAN MARGIN SEDIMENTS
Coastal plain
Sea level Shelf
Break
Margin sediments
Slope Rise
Continental crust Oceanic crust
100 km (A)
Sea level
Shelf Slope Accretionary prism
Volcanics
Oceanic crust Continental crust 100 km
Mantle
(B) Figure 2 Generalized cross-sections of (A) divergent and (B) convergent-type continental margins. Note the different scales. The divergent margin shows thick sedimentary deposits that have prograded into the ocean basin. Margin deposits in the convergent setting are much thinner, and most deposition occurs below the shelf break in the accretionary prism.
sediment transport is intercepted by trenches and elevated topography that inhibit the long-distance movement of turbidity currents. Seaward of this, though, there is frequently a wedge of deformed marine sediments called an accretionary prism that derives from material scraped from the subducting ocean crust. This situation is often found along the eastern North Pacific margin. Along margins with major trench systems such as those of the western Pacific, the rise and accretionary prism may be altogether absent.
Margin Sediment Sources, Character, and Distribution Ocean margin sediments generally comprise small particles ranging in size from fine sands (62–125 mm) to silts (4–62 mm) and clays (o4 mm). Larger sediments are occasionally found in ocean margin deposits, including coarse sands to boulders, but these particles are generally either relict (not actively transported), of glacial or ice-rafted origin, or biogenically precipitated (i.e. shell or coral). Overall patterns of grain-size distribution in margin sediments depend on the regional climate, geology, and sediment transport processes operating on the margin.
Three major mineral groups comprise ocean margin sediments, including siliciclastics, carbonates, and evaporites, or any combination of these. Worldwide, however, siliciclastics dominate ocean margin sediments. Siliciclastics are almost exclusively silicon-bearing minerals that derive from the physical and chemical weathering of continental rocks. Because most of the hundreds of rock-forming minerals are weathered before reaching the coast, margin siliciclastics dominantly comprise quartz and four clay species: kaolinite, chlorite, illite, and smectite. There is often a small component (several percent) of remnant feldspars, micas, and heavy minerals (sp. gr. 42.85) as well. Quartz, although an abundant and ubiquitous mineral, is often useful for interpreting margin sediments because different grain sizes reflect varying sources (e.g. eolian input) or energy regimes. For clay minerals, though, relative abundance of the major species in part reflects regional climate, thus making them a useful parameter for interpreting terrestrial signals preserved on the margin (see Clay Mineralogy). Furthermore, within the heavy mineral fraction, more unique silicate species can help determine the specific source area (provenance) of margin sediments. This approach is often limited because of regionally common
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OCEAN MARGIN SEDIMENTS
mineralogy, but researchers have begun to use isotope ratios to target more precisely margin sediment source areas. Since ocean margin sediments and deposits are closely tied to continental fluxes, such detailed studies are critical for an integrated understanding of the terrestrial-margin–marine sedimentary system. Although nearly all siliciclastic margin sediments are derived from the continents, they may be delivered via different mechanisms such as glaciers, rivers, coastal erosion, and wind. Although the latter mechanism is important to deep-sea sedimentation, the others dominate transport to the margin. At high latitudes, glaciers are a major and sometimes sole sediment source. The characteristics of glacial margin sediments is highly variable because of the capacity of ice to transport sediments of all sizes, but in general glacial sediments are coarser and more poorly sorted near the ice front (e.g., till) and become progressively finer and better sorted with distance (e.g., proglacial clays). Examples of glacially dominated modern margins include south-east Alaska, Greenland, and Antarctica. Along most of the world’s margins, rivers dominate sediment delivery to the shelf. The grain size of river sediments is typical of margin deposits, mostly including fine sand, silt, and clay-sized particles. However, the range of textures and the total amount of sediment delivered to the margin varies with many factors, including the size, elevation, and climate of the river basin. Because rivers are largely point sources, coastal processes such as waves and tides are important controls on redistributing fluvial sediments once they reach the ocean margin. In contrast to siliciclastics, carbonate sediments are largely biogenic in origin and are precipitated in situ by various corals, algae, bryozoans, mollusks, barnacles, and serpulid worms. Carbonates can be a dominant or significant component of margin sediments where the flux of siliciclastics is not great, often in arid regions or on the outer shelf. Carbonates do not usually form on the slope because of unstable bottom surfaces and depths below the photic zone. In general the production of carbonate sediments is higher in warmer low-latitude waters, but they are found along margins throughout the world. Unlike siliciclastic sediments, carbonates are not widely transported and tend to be concentrated within local regions on the shelf, where the build-up of bioherms and reefs may comprise locally important structural features. Overall, areas of purely carbonate production are rare and limited to ocean margins detached from the continents, such as the Bahamas. In contrast, mixed siliciclastic–carbonate margins are not uncommon and include regions such
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as the Yucatan and Florida peninsulas, Great Barrier Reef, and Indonesian islands. In very arid regions, margins along smaller marine basins may also support evaporite deposits. These minerals are precipitated directly from the water column to form thick salt deposits. Evaporites are also plastic, low-density sediments that frequently migrate upwards to form salt diapirs, causing major deformation of the overlying margin strata. Evaporite margin deposits are presently forming in the Red Sea and are also an important component of ancient margin sequences around the Mediterranean and Gulf of Mexico. Along many modern margins, relict and recently eroded sediments comprise the bulk of material that is actively transported on the shelf. An example of a relict margin is the US East coast, where modern fluvial sediment is trapped in coastal estuaries and embayments, thus starving the shelf of significant input. Thus, the shelf there is characterized by a thin sheet of reworked sands (remnant from the Holocene transgression) that overlie much older, partly indurated deposits. Along mountainous high-energy coastlines where rivers are absent or very small, the erosion of headlands can be a major source of sediment to the margin. Often these are also convergent margins with steep shelves that aid in advecting eroded near-shore sediments to the outer shelf and slope. Examples include portions of the eastern Pacific margin and south-eastern Australia (divergent).
Margin Sediment Transport Rivers are one of the major sources of sediment delivered to the shelf, and thus the fate of river plumes is an important control on the distribution of ocean margin sediments. The initial dispersal of the river plume and the subsequent sediment deposition are controlled by waves, tides, and three basic effluent properties that include the inertia and buoyancy of the plume, and its frictional interaction with the seafloor. In general, wave energy has the effect of keeping sediments close to shore, particularly sands that may be reworked onshore and/or transported alongshore by wave-driven currents. Margins that are wave dominated are typically steep and narrow and found in the high-latitudes, such as southern Australia, southern Africa and the margins of the north and south Pacific (i.e., high wind-stress regions). In contrast to waves, tides force a general onshore–offshore movement of sediment, most significantly along the coast where tidal energy is focused. Tide-dominated margin sediments typically occur on wide shelves and in shallow marginal basins such as the Yellow or North Seas. Compared with
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OCEAN MARGIN SEDIMENTS
waves and tides, the effect of river plume dynamics on sediment dispersal is more varied, particularly as complex feedbacks exist between coastal geology and riverine, wave, and tidal processes. The fate of sediments along margins is not a simple process of transport, deposition, and burial, but rather involves multiple cycles of erosion, transport, and deposition before being preserved and incorporated into ocean-margin strata (Figure 3). Thus, the triad of processes discussed above largely controls the initial phase of dispersal and deposition at the coast and inner shelf. The implication is that sediments undergo a succession of transport steps prior to preservation, and between each step the controlling processes, and thus sediment character, progressively change from riverine to coastal, shelf, and marine-dominated signals. Thus, beyond the coastal zone a different set of mechanisms is responsible for advecting sediments across the margin to the outer shelf, slope, and rise. On the shelf where seafloor gradients are often too low for failure (i.e., mass-wasting events), storms can be a major agent for cross-shelf transport. Two components of storms are involved in this process. First, strong winds generate steep, long-period waves that can resuspend sediment at much greater depths than fair-weather waves (e.g., a storm wave with a 12 s period will begin to affect the seafloor at B110 m depth). Second, resuspended sediments may be transported offshore via a near-bottom return flow (set up by wind and wave stresses) or by group-bound infragravity (long period) waves. Although these mechanisms for seaward sediment transport are well recognized, their magnitude, frequency, and distribution are not precisely known.
A second and well-described mode for sediment transport across ocean margins is the suite of gravitydriven mass movements that includes slumps, slides, flows, and currents (Figure 4). This suite comprises nearly a continuum of properties from the movement of consolidated sediment blocks as slumps and slides to the superviscous non-Newtonian fluids of debris and sediment flows to the low-density suspended load of turbidity currents. Mass wasting events are common, widespread, and most frequently occur where slopes are oversteepened and/or sediments are underconsolidated (i.e., low shear strength). These last two factors are related since the threshold for oversteepening is partly a function of the sediment’s shear strength, with possible failure occurring at gradients o11 in poorly consolidated sediments (e.g., 80–90% water content) and relative stability found on slopes of 201 for well-consolidated material (e.g., 20–30% water content). Many other factors also contribute to mass movements including groundwater pressure, sub-bottom gas production, diapirism, fault planes, and other zones of weakness. In addition to these intrinsic sedimentary characteristics, trigger mechanisms are an important control on mass movements. Triggers may include earthquakes, storms, and wave pumping, each of which may have the effect of inducing shear along a plane of weakness or disrupting the cohesion between sediment grains (i.e., liquefaction). Where trigger mechanisms are frequent and intense, such as along a tectonically active or high-energy margin, mass wasting may occur in deposits that would be stable in a passive, low-energy setting. Although storm activity can be important on the shelf, mass movements are by far the most dominant
Estuary
Wind-driven alongshelf flows
ach
Be
ps
Ri
Ve diff r tical usi on
ar
Estuarine plumes
Front
s ve Wa
Internal waves
pth De ) (m
s ave ty w i v ra rag Inf
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tor
20 Physical and biologic mixing of sediment column
Botto
m bo u layer ndary
Steep gradients in cross-shelf sediment flux
Inner shelf
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an
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en
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ys nb
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Mid shelf
Figure 3 Conceptual diagram illustrating the major physical processes responsible for transporting sediment across the continental shelf. Note the many processes that contribute to the resuspension, erosion, and transport of sediment in this region. (Reproduced with permission from Nittrouer and Wright, 1994.)
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OCEAN MARGIN SEDIMENTS
Mud diapirs and growth faults (creep)
Shelf, delta platform (High sedimentation rate)
Mud flows with overlapping toe lobes
Shelf break erosion Sediment Wind blown accumulation in canyon sand head
Slide
Soft and semiconsolidated slope sediments, mainly slumps and mud flows
Slope valley (feeder channel for delta fan) Slump
Active faults
(strong relief)
Young sediments
Older rocks
Deep-sea delta fan Large active and buried sand-filled channel systems, mainly mud turbidites, some mud flows
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Canyon deep-sea fan system Upper fan: some slumps, flows of grains, debris and mud, channel fills. Lower fan: mainly sand turbidites, some channel fills
Slope apron Slumps, mud flows predominantly mud turbidites
Transition to high- and low-density turbidity currents, sand and mud turbidites
Rockfall, slumps, and debris flows (polymict), undercutting of slope by contour currents
Figure 4 Summary diagram of sedimentary features and transport mechanisms typical of continental margins. Note the variety of mass-wasting events that originate on the slope and extend onto the rise. (Reproduced with permission from Einsele G (1991) Submarine mass flows, deposits and turbidities. In: Einsele G et al. (eds) (1991) Cycles and Events in Stratigraphy. New York: Springer-Verlag.)
mechanism for transferring sediment to the slope and rise. Indeed, estimates indicate that up to 90% of continental rise deposits derive from turbidity currents. In terms of their occurrence, slumps (movement of consolidated sedimentary blocks) are most common along steep gradients such as the continental slope or along the channels and walls of submarine canyons. Slumps and slides may move over distances of meters to tens of kilometers, and at rates of centimeters per year to centimeters per second. Although slumps and slides do not frequently travel great distances, these energetic events often generate turbidity currents at their front, whic
