Biogeochemistry of Marine Dissolved Organic Matter

Citation preview

Biogeochemistry of Marine Dissolved Organic Matter

This Page Intentionally Left Blank

Biogeochemistry of Marine Dissolved Organic Matter Edited by

Dennis A. Hansell University of Miami Rosenstiel School of Marine and Atmospheric Science Miami, Florida

Craig A. Carlson University of Californian Santa Barbara Santa Barbara, California

/ ^ ACADEMIC PRESS V — ^ An Elsevier Science Imprint Amsterdam Boston London New York Oxford Paris San Diego San Francisco Singapore Sydney Tokyo

This book is printed on acid-free paper. ®

Copyright © 2002 by Elsevier Science (USA) All Rights Reserved. No part of this publication may be reproduced or transmitted in any form or by any means, electronic or mechanical, including photocopy, recording, or any information storage and retrieval system, without permission in writing from the publisher. Requests for permission to make copies of any part of the work should be mailed to: Permissions Department, Academic Press, 6277 Sea Harbor Drive, Oriando, Florida 32887-6777 Academic Press An Imprint of Elsevier of Elsevier Science 525 B Street, Suite 1900, San Diego, California 92101-4495, USA http://www.academicpress.com Academic Press 32 Jamestown Road, London NWl 7BY, UK http://www.academicpress.com Library of Congress Catalog Card Number: 2001096950 International Standard Book Number: 0-12-323841-2 PRINTED IN THE UNITED STATES OF AMERICA 02 03 04 05 06 07 MM 9 8 7 6 5 4

3 2 1

For the support and balance that only family can provide we dedicate this hook to our beloved spouses Paula and Alison^ and our children Allison and Rachel, and Matthew and Hayden.

This Page Intentionally Left Blank

Contents

Contributors Foreword Preface

xiii xv xxi

Chapter 1

Why Dissolved Organics Matter? John I. Hedges I. 11. III. IV. V. VI. VII.

Introduction 1 DOM Research Pre-1970 2 DOM Research in the 1970s 7 DOM Research in the 1980s 11 "New" DON and DOC 13 Why Dissolved Organics Matter 23 What did we Learn? 25 References 27

Chapter 2

Analytical Methods for Total DOM Pools Jonathan H. Sharp I. II. III. IV.

Introduction 35 Dissolved Organic Carbon Analysis Dissolved Organic Nitrogen Analysis Dissolved Organic Phosphorus Analysis

37 45 49

Contents

V. Multielemental Methods 51 VI. TheLimitsof Elemental Analyses 51 VII. The Need for Continual use of Reference Materials References 54

52

Chapter 3

Chemical Composition and Reactivity Ronald Benner

I. Introduction 59 II. Distribution and Chemical Characteristics of Bulk Marine DOM 64 III. Major Topics of Ongoing and Future Research About the CycUng of DOM 80 References 85 Chapter 4

Production and Removal Processes Craig A. Carlson

I. II. III. IV. V. VI.

Introduction 91 DOM Production Processes 92 DOM Removal Processes 116 DOM LabiUty 123 DOM Accumulation 133 Summary 137 References 139

Chapter 5

Dynamics of DON Deborah A. Bronk

I. II. III. IV. V.

Introduction 153 Concentration and Composition of the DON Pool Sources of DON 186 Sinks for DON 207 DON l\imover Times 226

154

ix

Contents VI. Summary References

227 231

Chapter 6

Dynamics of DOP D. M. Karl and K. M. Bjorkman

I. Introduction 250 II. Terms, Definitions, and Concentration Units 253 III. TheEarly Years of Pelagic Marine P-Cycle Research (1884-1955) 258 IV. The Pelagic Marine P-Cycle: Key Pools and Processes V. Sampling, Incubation, Storage, and Analytical Considerations 266 VI. DOP in the Sea: Variations in Space 280 VII. DOP in the Sea: Variations in Time 294 VIII. DOP Pool Characterization 306 IX. DOP Production, Utilization, and Remineralization X. Conclusions and Prospectus 347 References 348

262

334

Chapter 7

Marine Colloids and Trace Metals Mark L Wells

I. II. III. IV. V. VI. VII. VIII. IX. X.

Introduction 367 Definition of Marine Colloids 369 Analytical Methods 372 Metal Content of Marine Colloidal Matter 380 The Chemical Form of Colloidal Metals 385 Particulate-Based Estimates of Colloidal Metal Concentrations 388 Sources of Metal-Complexing Colloidal Ligands 389 Measurement of Colloid Reaction Rates 390 The Biological Availability of Colloidal Bioactive Metals Summary 396 References 397

395

Contents

Chapter 8

Carbon Isotopic Composition of DOM James E. Bauer

I. Introduction 405 II. Conventions and Definitions for Expressing Isotopic Contents of DOC 407 III. Methods for Extracting DOC from Seawater for Isotopic Analysis 413 rV. Measurements and Distributions of 6^^C and A^'^C in Marine DOC 415 V. Applications of ^^^C and A^^C in Marine DOC Cycling Studies 430 VI. Summary and Future Challenges 443 References 446 Chapter 9

Photochemistry and the Cycling of Carbon, Sulfur, Nitrogen and Phosphorus Kenneth Mopper and David J. Kieber

I. Introduction 456 II. Photochemical Transformation of Riverine and Marsh-Derived DOM Inputs to the Sea 457 III. Impact of Photochemistry on Elemental Cycles 458 IV. Unresolved Questions and Future Research 476

References Appendix 1 Appendix 2 Appendix 3 Appendix 4

479 490 498 500 503

Chapter 10

Chromophoric DOM in the Coastal Environment Neil V. Blough and Rossana Del Vecchio

I. Introduction 509 II. Optical Properties 513 i n . Distribution 532

xi

Contents IV. Sources and Sinks 534 V. Summary and Future Areas of Research References 540

539

Chapter 11

Chromophoric DOM in the Open Ocean Norman B. Nelson and David A. Siegel I. II. III. IV. V.

Introduction 547 Characterization of CDOM 549 Observed CDOM Dynamics 557 Global CDOM Distribution Patterns 561 Relationship Between DOM and CDOM in the Open Ocean 567 VI. Implications for Photochemistry and Photobiology VII. Needs for Future Advances 571 References 573

568

Chapter 12

DOM in the Coastal Zone Gustave Cauzvet I. II. III. IV.

Introduction 579 River Inputs 580 Estuarine Processes 588 Accumulation of DOM in the Coastal Zone and Export Processes 595 V. Conclusions 600 References 602

Chapter 13

Sediment Pore Waters David J. Burdige I. II. III. IV.

Introduction 612 Dissolved Organic Carbon in Sediment Pore Waters Dissolved Organic Nitrogen (DON) 631 DOM Compositional Data 636

614

Contents

V. The Role of Benthic DOM Fluxes in the Ocean Carbon and Nitrogen Cycles 641 VI. The Role of Pore-Water DOM in Sediment Carbon Preservation 648 VII. Conclusions and Suggestions for Future Research 650 Appendix: A Description of the DOM Advection/Diffusion/ Reaction Model 651 References 653 Chapter 14

DOC in the Arctic Ocean LeifG. Anderson

I. Introduction 665 II. Sources of DOC to the Arctic Ocean 667 III. Composition and Distribution of DOC within the Arctic Ocean 674 IV. Summary of Sources and Sinks 679 References 681

Chapter 15

DOC in the Global Ocean Carbon Cycle Dennis A. Hansell

I. II. III. IV. V. VI.

Introduction 685 Distribution of DOC 687 Net Community Production of DOC 697 Contribution of DOC to the Biological Pump Research Priorities 709 Summary 711 References 711

Chapter 16

Modeling DOM Biogeochemistry James R. Christian and Thomas R. Anderson

I. Introduction 717 II. Ecosystem Modeling Studies

719

702

Contents III. Modeling the Role of DOM in Ocean Biogeochemistry IV. Discussion and Conclusions 743 References 747

Index

734

757

Contributors

Numbers in parentheses indicate page numbers on which the authors contributions begin.

Leif G. Anderson (665), Analytical and Marine Chemistry, Goteborg University Goteborg, Sweden Thomas R. Anderson (717), George Deacon Division, Southampton Oceanography Centre, Southampton United Kingdom James E. Bauer (405), School of Marine Science, College of William and Mary, Gloucester Point, Virginia

xiv

Contributors

Ronald Benner (59), Department of Biological Sciences and Marine Science Program, University of South Carolina, Columbia, South Carolina Neil V. Blough and Rossana Del Vecchio (509), Department of Chemistry and Biochemistry, University of Maryland College Park, Maryland Deborah A. Bronk (153), Virginia Institute of Marine Science, College of WilHam and Mary, Gloucester Point, Virginia David J. Burdige (611), Department of Ocean, Earth, and Atmospheric Sciences, Old Dominion University Norfolk, Virginia Craig A. Carlson (91), University of California, Santa Barbara, Department of Ecology, Evolution and Marine Biology, Santa Barbara, California Gustave Cauwet (579), Laboratoire d'Oceanographie Biologique (UMR CNRS 7621), Observatoire Oceanologique, Banyuls sur mer, France James R. Christian (717), Universities Space Research Association, NASA Goddard Space Flight Center, Code 970.2 Greenbelt, Maryland Dennis A. Hansell (685), University of Miami, Division of Marine and Atmospheric Chemistry, Rosenstiel School of Marine and Atmospheric Science, Miami, Florida John I. Hedges (1), School of Oceanography, University of Washington, Seattle, Washington D. M. Karl and K. M. Bjorkman (249), Department of Oceanography, School of Ocean and Earth Science and Technology, University of Hawaii Honolulu, Hawaii David J. Kieber (455), College of Environmental Science and Forestry Chemistry Department, State University of New York Syracuse, New York Kenneth Mopper (455), Department of Chemistry and Biochemistry, Old Dominion University Norfolk, Virginia Norman B. Nelson and David A. Siegel (547), Institute for Computational Earth System Science, University of California, Santa Barbara, Santa Barbara, California Jonathan H. Sharp (35), Graduate College of Marine Studies, University of Delaware, Lewes, Delaware Mark L. Wells (367), School of Marine Sciences, University of Maine, Orono, Maine

Foreword

Few of us really have intuitive concepts of the differences among ocean ecosystems. Ecosystems on land clearly look different from one another - contrast, for example, the outward appearances of deserts and savannas. Yet oligotrophic gyres and continental shelves, the oceanic analogs of these terrestrial systems, look nearly identical to the unaided eye, and we have to look more deeply (sometimes literally) to perceive the differences. Nearly all terrestrial ecosystems rest, physically and functionally, on an organic-rich soil foundation. Dissolved organic matter (DOM) is the soil of the sea - a large, biochemically resistant reservoir of organic matter providing a substrate for life, and a source for nutrient regeneration, ion exchange capacity, light and heat absorption, and so on. Marine DOM, however, is much less conspicuous than terrestrial soil. It is, in fact, nearly invisible. In this book, Hansen and Carlson and the many contributing authors tell the story of making DOM, the soil of the sea, visible. Recently I was asked to provide a list to the International Geosphere-Biosphere Program (IGBP) of the top accomplishments and failures of the Joint Global Ocean Flux Study (JGOFS). I polled hundreds of scientists and students accessible via US JGOFS' e-mail lists and received numerous opinions about both the program's successes and failures. Interestingly, and as syndicated colunmist Dave Barry would say, "I am not making this up," one topic was on both lists - dissolved organic carbon, DOC! This book attests to the success of DOM studies in JGOFS (including carbon, nitrogen and phosphorus), and throughout ocean biogeochemistry over the past decade. Was it also a story of failure? The question is provocative and I want to explore it here, at least briefly. DOM has a long and distinguished history in marine chemistry and biology, dating to the early controversy as to whether or not this apparently large reservoir of organic matter was an important source of nutrition for marine animals (Krogh, 1934; Jorgensen, 1976). Duursma's (1963) monograph on the seasonal dynamics of DOC in the North Sea and North Atlantic revealed that the pool was an active and variable component of the marine ecosystem. The first radiocarbon dating of XV

xvi

Foreword

DOC by Williams et al. (1969) indicated that the vast majority of this globally significant carbon pool was long-lived and refractory - in both the deep as well as surface oceans. By the late 1980's, as JGOFS began to focus on properties of the ocean carbon system, DOC was perceived as uninteresting -just a large, inert pool without much discemable vertical structure or horizontal gradients. I recall Peter LeB. Williams showing me the DOC analyzer he developed. "Here's the world's best instrument for analyzing the ocean's most boring property!", he said. Added to this was controversy over the best analytical approach to quantify the bulk pool, which went back to Krogh and Keys (1934). Given this backdrop, the seminal paper on DOC analysis by Sugimura and Suzuki (1988) was greeted with great surprise and excitement. In demonstrating a new analytical method and some of its early results, they presented oceanic DOC profiles with surface gradients of several lOO's of /JLM and overall very much higher concentrations than revealed by earlier approaches. These findings made DOC interesting in several ways. Marine chemists seeking improvements to the thermodynamic description of the carbonate system in seawater saw in DOC a potential source of additional protolytes (Bradshaw and Brewer, 1988). Peter Brewer, the new Chair of U.S. JGOFS, was particularly energetic in advancing Suzuki's method and a newly recognized role for DOC in the carbon cycle. Perhaps the greatest push for the new, high DOC levels came from modelers. The 3-dimensional ocean modeling community became very interested in a DOM pool that had a longer lifetime than sedimenting particles and could be transported horizontally for long distances. In this behavior they saw the possible answer to the problem of nutrient trapping in models of the equatorial Pacific Ocean. Ray Najjar modeled DOM export to address the problem in his Ph.D. thesis (Najjar et al., 1992). Robbie Toggweiler discussed other aspects of high DOC levels in a still widely cited paper (Toggweiler, 1989). It was clear that a large and influential segment of the ocean community was prepared to embrace these exciting results. Suzuki's results led to upward revisions of the oceanic DOC inventory, and to an explosion of research on marine DOM, its chemistry, analysis and ecology. Yoshimi Suzuki became an overnight celebrity. He participated in the U.S. JGOFS North Atlantic Bloom Experiment, and measured DOC in May 1989 in close conjunction with Ed Peltzer from Brewer's lab at WHOI, again demonstrating high concentrations and spectacular variations in space and time. Perplexingly, there were no known biological processes to maintain variations in euphotic zone DOC stocks of about 1 mole C as found over scales of a few days or a few km. Yet his analyses made on the same cruise established one of the first direct estimates of DOC utilization by bacteria, and resulted in an influential estimate of bacterial growth efficiency (Kirchman et al., 1991). U.S. JGOFS sponsored two workshops, including a "bake-off" (alluding to high-temperature combustion techniques) to validate Suzuki's method (Williams, 1991). Although large segments of the conmiunity wanted the new results to be true, many marine

Foreword

xvii

chemists remained very skeptical. Reporting on the workshop results, Peter M. Williams reported, "Most strikingly, the ranges of variation in the mean DOC concentrations of the same water samples by the same types of DOC analyzer were almost as great as the entire data set... The DOC data from the different seawater analyses plot along three roughly parallel lines until reaching the high extreme of the measured range... and thus do not vary randomly. One explanation for this pattern is that analyses made by different instruments include blanks of varying magnitude." (Williams, 1991, p. 11).

Williams had it right, as was later demonstrated by Benner and Strom (1993) in the special issue of Marine Chemistry reporting the scientific results of the 1991 bake-off workshop. High-temperature, catalytic oxidation techniques for DOC analysis suffered from high instrument blanks that were not easily evaluated or corrected, leading to variable and high offsets in apparent DOC concentrations. In the meantime, Eiichiro Tanoue measured DOC in the same region of the northwestern Pacific assessed earlier by Sugimura and Suzuki (1988), finding much lower concentrations and less pronounced vertical gradients (Tanoue, 1992). In response to these new findings, Suzuki began a reassessment and reanalysis of his original results. In a statement of extraordinary courage and grace he retracted the results that had caused so much excitement (Suzuki, 1993; see also Hedges et al., 1993). Thus, we see in this series of events a scenario familiar in the history of science. An idea, stimulated by technological innovation, was advanced and tested. Great excitement ensued and the new results suggested new solutions to recognized problems. More scientists saw a subject in a new way. But with increased scrutiny, the method was found wanting and the results were ultimately rejected. I think this is the reason some scientists have tended to regard oceanic DOC measurement as a failure... the initial results didn't hold up. To some, Suzuki is the villain of the story, too quick to accept apparently spectacular results without adequate testing. I view the situation differently. As a result of the excitement generated by the original paper, and by Brewer's and others' strong advocacy of it, many others began to think in new ways about DOM in the sea. They wrote proposals and started new research. The technical aspects of DOC analysis were examined in an unprecedented manner, resulting in new instruments with great precision, capable of resolving 1 /xM differences in DOC concentration. There is today a recognized DOC analytical standard. These developments made possible direct detection of bacterial utilization of the bulk DOC pool, thus allowing us to assess the varying lability of the bulk DOM pool, insights expanded upon the results of Barber (1968) and Ogura (1972) a generation earher. Following the idea pursued t>y Najjar and colleagues, DOC eventually became recognized as an important vector of export production (Copin-Montegut and Avril, 1993; Carlson et al., 1994). Increased precision enabled detection of deep-ocean DOC concentration gradients and basin-scale differences in DOC (Hansell and Carlson, 1998), opening its use

xviii

Foreword

as a new geochemical tracer. Although the NABE study lacked reliable DOC data, all subsequent JGOFS studies had successful DOC research components. Oceanic DOM is now recognized as an important component of the biogeochemical system and possibly a barometer of global change (Church et al., 2002). Most importantly, we can today regard marine DOC as a dynamic component in the global carbon cycle. Success or failure? Read this book and be the judge. Hugh W. Ducklow School of Marine Science The College of William and Mary

REFERENCES Barber, R. T. (1968). Dissolved organic carbon from deep waters resists microbial oxidation. Nature 220,274-5. Benner, R. and Strom, M. (1993). A critical evaluation of the analytical blank associated with DOC measurements by high-temperature catalytic oxidation. Mar. Chem. 41,153-60. Bradshaw, A. L. and Brewer, R G. (1988). High precision measurements of alkalinity and total carbon dioxide in seawater by potentiometric titration. 1. Presence of unknown protolyte(s)? Mar. Chem. 23,69-86. Carlson, C. A., Ducklow, H. W. and Michaels, A. F. (1994). Annual flux of dissolved organic carbon from the euphotic zone in the northwestern Sargasso Sea. Nature 371,405^08. Church, M. J., Ducklow, H. W. and Karl, D. M. (2002). Multi-year increases in dissolved organic matter inventories at Station ALOHA in the North Pacific Subtropical Gyre. Limnol. Oceanogr. 47,1-10. Copin-Montegut, G. and Avril, B. (1993). Vertical distribution and temporal variation of dissolved organic carbon in the northwestern Mediterranean Sea. Deep Sea Res. 40, 1963-1972. Duursma, E. K. (1963). The production of dissolved organic matter in the sea, as related to the primary gross production of organic matter. Netherlands Journal of Sea Research 2, 85-94. Hansen, D. A. and Carlson, C. A. (1998). Deep ocean gradients in dissolved organic carbon concentrations. Nature 395, 263-266. Hedges, J., Lee, C. and Wangersky, P. J. (1993). Conmients from the editors on the Suzuki statement. Mar Chem. 41, 289-290. Krogh, A. (1934). Conditions of life in the ocean. Ecol. Monogr 4,421^29. Krogh, A. and Keys, A. B. (1934). Methods for the determination of dissolved organic carbon and nitrogen in sea water. Biol. Bull. 67,132-144. J0rgensen, C. B. (1976). August Putter, August Krogh and modem ideas on the use of dissolved organic matter in the aquatic environment. Biol. Rev. 51, 291-308. Kirchman, D. L., Suzuki, Y., Garside, C. and Ducklow, H. W. (1991). High turnover rates of dissolved organic carbon during a spring phytoplankton bloom. Nature 352,612-^. Najjar, R. G., Sarmiento, J. L. and Toggweiler, J. R. (1992). Downward transport and fate of organic matter in the ocean: simulations with a general ocean circulation model. Global Biogeochem. Cycles 6,45-76. Ogura, N. (1972). Rate and extent of decomposition of dissolved organic matter in the surface water. Mar. Biol. 13, 89-93. Sugimura, Y. and Suzuki, Y. (1988). A high-temperature catalytic oxidation method of non-volatile dissolved organic carbon in seawater by direct injection of liquid samples. Mar. Chem. 14,105-131.

Foreword

xix

Suzuki, Y. (1993). On the measurement of DOC and DON in seawater. Mar. Chem. 41, 287-288. Tanoue, E. (1992). Vertical distribution of dissolved organic carbon in the North Pacific as determined by the high temperature catalytic oxidation method. Earth Planet. Sci. Lett. I l l , 201-216. Toggweiler, J. R. (1989). Is the downward dissolved organic matter (DOM) flux important in carbon transport?, In "Productivity of the oceans: present and past" (W. H. Berger, V. S. Smetacek and G. Wefer, Eds.), pp. 65-83, Wiley. Williams, P. M., Oeschger, H. and Kinney, P. (1969). Natural radiocarbon activity of the dissolved organic carbon in the northeast Pacific Ocean. Nature 224,256-258. Williams, P. M. (1991). Scientists and industry reps attend workshop on measuring DOC in natural waters. US JGOFS News 3(1), 1,5,11.

This Page Intentionally Left Blank

Preface

Efforts by the ocean science community to understand the cycHng of the major bioactive elements (C, N, P) in the ocean expanded rapidly in the last decade and continues today. The intensive focus on elemental cycling resulted from society's need to determine the role of the ocean in global cHmate change. By the beginning of the 1990's, the fundamentals of the biological processes involved in the transformations of the major elements were identified. The next phase of research required linking the biological processes to the very large oceanic reservoirs of the major elements. Establishing this linkage between processes and reservoirs falls into the discipline of biogeochemistry. One of the Earth's largest bioactive reservoirs of carbon is dissolved organic matter (DOM) in the ocean. With a stock of 700 Pg C in the global ocean, the pool is approximately equal in size to the stock of carbon resident in atmospheric CO2. Prior to the 1990's, this major pool of carbon was primarily evaluated from a geochemical perspective; resolving the composition of the pool was a central goal. With the onset of an enhanced biogeochemical perspective of nutrient cycling, the scientific questions began to rest broadly on the role of DOM in the oceanic C, N and P cycles. To determine the function of DOM in the elemental cycles, vast intellectual and financial capital was expended throughout the 1990's. Central questions were: can we accurately, with community wide consistency, measure the concentrations of dissolved organic matter in the ocean; what are the distributions of the dissolved organic C/N/P pools and what processes controls these distributions; what are the rates, biogeographical locations and controls on elemental cycling through the pools; what are the biological and physicochemical sources and sinks; what is the composition of the pools and what does this tell us about elemental cycling? Finally, do we understand DOM in elemental cycling well enough to accurately represent the processes in numerical models? In this book, the progress of the last decade in answering these questions is reported and synthesized by key contributors to those advances. The book opens with a chapter by J. Hedges, providing historical perspective for the work of XXI

xxii

Preface

the 1990's, as well as context for the succeeding chapters. An important obstacle that had to be breached before significant biogeochemical advances could be made was coordinated improvement in the methods for determining the bulk DOM pool concentrations. J. Sharp, a leader in those community efforts, reviews methodological advances in Chapter 2. Study of the chemical and isotopic compositions of DOM has provided unique information on elemental cycling. This work is reviewed in chapters by R. Benner and J. Bauer. The biological cycling of the major elements (C, N, P) through DOM is reviewed in chapters by C. Carlson, D. Karl and K. Bjorkman, and D. Bronk. Particular emphasis is placed in these chapters on marine microbes as active agents in the processing of DOM. Colloidal organic matter, with special focus on interactions with metals, is covered by M. Wells. Photochemical reactivity of DOM, and implications for elemental cycling, is discussed by K. Mopper and D. Kieber. The contribution of optically active (chromophoric) DOM in bio-optical processes is covered by N. Blough and R. Del Vecchio for the coastal ocean and by N. Nelson and D. Siegel in the open ocean. The role of DOM in the ocean margins and interfaces (i.e., the coastal realm, the sediments, and the Arctic Ocean) is reviewed in chapters by G. Cauwet, D. Burdige, and L. Anderson, respectively. A review of the global ocean distribution and broad scale transformations of DOM is presented by D. Hansell. The book closes with discussion on the advances for the inclusion of DOM in both ecosystem and global circulation models by J. Christian and T. Anderson. Many scientists in the ocean science conmiunity have developed a strong biogeochemical view of the ocean. This book provides a firm foundation for their forays into the biogeochemistry of marine organic matter. The book maintains a particular focus on DOM in elemental cycling, and therefore does not revisit the many, well-documented advances made in organic geochemistry during the previous decades. Attention is paid largely to the marine environment, with coverage of fresh water systems only at its interface with the marine realm. The book is directed at professional ocean scientists and advanced students of biological and chemical oceanography. Many individuals and organizations must be thanked for support of the science that provided content for this book, as well as to development of the book itself. The U.S. federal agencies supporting much of what has been reported here, including individual research by the chapter authors, are the National Science Foundation, the National Oceanographic and Atmospheric Administration, and the National Aeronautics and Space Administration. The agency program managers who have provided invaluable support to we editors are Neil Anderson, Lisa Dilling, Don Rice, Phil Taylor, and Jim Todd. The U.S. JGOFS program, particularly the Scientific Steering Committee and the Planning Office, provided consistent support to ensure that our understanding of DOM in marine elemental cycles was advanced. Their vision and encouragement was necessary for the many advances reported in this book to be realized. D. A. Hansell and C. A. Carlson

Chapter 1

Why Dissolved Organics Matter John I. Hedges School of Oceanography, University of Washington, Seattle, Washington

I. II. III. IV. V.

Introduction DOM Research Pre-1970 DOM Research in the 1970s DOM Research in the 1980s "New" DON and DOC

VI. Why Dissolved Organics Matter VII. What did we Learn? References

I. INTRODUCTION As this book attests, research on dissolved organic matter (DOM) in seawater has burgeoned in the past decade. This increase in activity is evident not only from the growing number of articles published each year in the scientific literature, but also from the topical breadth and broad integration of present research. The oceanographic community's perception of DOM has evolved from an emphasis on a dilute and largely separate pool of remarkably old and static substances to the current view of a dynamic assemblage of organic molecules that interact with each other, trace metals, and living organisms over a broad continuum of space and time scales. The sparingly reactive components of this molecular continuum that persist and change on time scales sampled by conventional oceanographic surveys represent a small molecular outcrop of a churning mass of molecules through which much of the total primary production of the ocean cycles. To better understand what is to come in this chapter and book, it is useful to keep in mind that investigations of DOM in seawater have followed two fundamentally Biogeochemistry of Marine Dissolved Organic Matter Copyright 2002, Elsevier Science (USA). All rights reserved.

1

2

John I. Hedges

different strategies. The first is a holistic approach focusing primarily on the total concentration, bulk properties, and collective behavior of the entire mixture of molecules that make up the operationally defined DOM pool. Examples would be measurements of the total dissolved organic carbon (DOC) or dissolved organic nitrogen (DON) concentrations, determinations of bulk spectral or isotopic compositions, and estimates of cumulative oxygen and nutrient changes attending microbial attack of the entire organic mixture. This strategy has the major advantage of yielding characteristics that are representative of the entire DOM pool, but the information obtained is typically limited and highly biased toward the less reactive components of the mixture that accumulate over time. In contrast, the reductionist approach has been to target selected fractions of the total mixture for detailed analyses of specific features that might then be meaningfully extrapolated back to the bulk pool. The most common form of reductionism is the chromatographic analysis of specific biochemical components of seawater DOM. This particular strategy can yield a wealth of information on structural features, stereochemistries, and reaction pathways and dynamics. However, molecular-level analyses are highly selective for individual biochemical classes (or subclasses), which in turn often comprise a tiny, and not necessarily representative, fraction of bulk DOM. Thus, major uncertainties arise in extrapolating from detailed molecular-level information to the whole DOM pool, and especially to its emergent properties. This introductory chapter emphasizes the oceanographic community's perceptions of the entire DOM pool from a bulk chemical perspective, bringing in biochemical and microbiological information primarily as it pertains to the larger view. While this focus on collective properties necessitates that substantial advances at the biochemical level will not be highlighted, it does allow better historic continuity and further development of broad issues pertaining to oceanography in general. This chapter recaps selected experimental and conceptual developments extending from the last century up through the Seattle DOC/DON Workshop Report (Hedges and Lee, 1993) that have led to the modem dynamic view of oceanic DOM presented in the following chapters.

11. DOM RESEARCH PRE-1970 By 1970, study of seawater DOM had already been under way for almost a century (see review by Kalle, 1966). Glass fiber or silver filters available in the mid-20th century had minimal pore sizes of ^0.45-1.0 /xm and became the basis of the operational definition that "dissolved" materials pass such filters whereas "particulate" matter does not (Fig. 1). This definition persists to today, although we now know that seawater contains a continuum of discrete units stretching from the size of whales to that of a water molecule, with no discemable break in abundance in the micrometer range (Sharp, 1973a). The traditional definition can be useful.

W/zy Dissolved Organics Matter mm Meters

I

urn

10-^I III I 10-^ mill III

nm

I

10-5 10-^ IN III nil, lllllilJI

10-^

Partlcuiate

10-8 iiiiiii III

I

I

lo-^ mill I

III

10-^°

Dissolved Colloids

^fel Sand

I [

Viruses

_

|

Macyomolecutes "[

'^rCliyJ Screen Sieves

^^ -^iijfi'lWr,.,L^ Papers f Ultrafilters

Figure 1

Molecular

-^

^-

Sieves

The continuum of sizes and separation methods for organic matter in seawater.

however, because particles smaller than 1/xm are not prone to sink (Duursma, 1961) and all living organisms other than viruses and small bacteria fall into the particulate fraction. Colloidal particles, constituting the upper size range (0.0011.0 jjim) of the DOM continuum, correspond in minimal size to approximately a six-sugar oligosaccharide (Fig. 1). Following several largely unsuccessful early attempts (e.g., Piitter, 1909; Raben, 1910) to quantify the dissolved organic contents of seawater, Krogh and Keys (1934) published comparatively reproducible methods for the determination of both DON and DOC in seawater. The DON method was based on a micro-Kjeldahl (sulfuric acid hydrolysis) procedure, whereas DOC was quantified (after chloride removal) by wet oxidation in aqueous chromic acid. Using these methods, Krogh (1934a) measured the first full water column profiles of DOC and DON in the open ocean off Bermuda. He found uniform concentrations of organic material from the surface down and concluded that seawater DOM is chronologically old, chemically and biochemically inert, and insignificant as a food source for organisms in the deep sea (Krogh, 1934b). The following year, however, Waksman and Carey (1935) demonstrated in a series of culturing experiments that bacteria decompose DOM from surface seawater in a matter of days, with attending increases in inorganic nitrogen and decreases in dissolved oxygen. Kalle (1937) used UV absorption to detect yellow organic substances in the waters of the North Sea and open North Atlantic. Although spectroscopically similar to DOM in rivers, seawater "gelbstoff" was recognized to have a predominant

4

John I. Hedges

marine origin (Kalle, 1949). Kalle (1949) also reported an organic component of seawater DOM that gives a bluefluorescencewhen irradiated with long-wavelength ultraviolet light and appears to have a predominantly terrestrial source. Early attempts to isolate seawater DOM by sorption onto charcoal (Wilson and Armstrong, 1952; Johnston, 1955) or extraction with nonpolar solvents (Slowley et al, 1959; Chanu, 1959) were successful, although subsequent chemical characterizations were primarily limited to demonstrating UV absorbance and the presence of trace amounts of fatty acids (Jeffrey and Hood, 1958). Various laboratory experiments (e.g., Fogg and Boalch, 1958) demonstrated that marine algae (especially phaeophyta) are potential direct sources of seawater DOM. At this time, amino acids and carbohydrates were known to spontaneously condense (although at elevated temperatures) to produce melanoidin polymers (Maillard, 1913) that exhibit many of the spectral qualities of marine DOM (Kalle, 1966). By the early 1960s, DOC was measured at concentrations on the order of 1 mg/L (83.3 /xM) and found to be more concentrated in surface ocean water than at depth (Kay, 1954; Plunkett and Rakestraw, 1955; Duursma, 1961). In addition, a variety of component biochemicals, including simple sugars, low-molecular-weight acids, and vitamin B12, had been detected in seawater (Vallentyne, 1957; Hood, 1970; Duursma, 1965). A wave of pioneering field studies during the 1960s served mainly to strengthen the perception of deep-ocean DOM as a largely static pool. Improved wet chemical oxidation methods for seawater DOM (e.g., the persulfate adaptation of Menzel and Vacarro, 1964) became the basis for extensive surveys of DOC concentrations in various oceans (e.g., Menzel, 1964; Menzel and Ryther, 1968). Menzel's 1964 study of DOC distributions in the western Indian Ocean was by far the most extensive to that time with respect to the number of stations (39) and depths (1-2000 m) sampled. In addition to synoptic temperature and salinity data for each sample, this study included ^"^C-based measurements of primary production under simulated euphotic zone conditions. No apparent correlation between DOC concentration and primary production rates was observed in surface ocean waters. Although DOC concentrations below 200 m ranged geographically between 0.2 and 2 mg/L ('^ 15-170 /xM), these gradients covaried linearly with salinity and thus appeared to result primarily from mixing of different water masses with characteristically different DOC signatures. Menzel (1964) concluded, "carbon in solution and in particulate form in the ocean is extremely stable and subject to limited change by biological activity." Menzel and Ryther (1968) soon published a more detailed study of dissolved and particulate organic carbon (POC) distributions in discrete water samples collected over the entire water column at 14 stations in the southern Atlantic Ocean. In contrast to the Indian Ocean survey, dissolved oxygen was directly measured for each sample, along with temperature and salinity. Relatively constant DOC concentrations (35 ± 5 /xM) were observed at depths greater than 500 m throughout the South Atlantic (Fig. 2). Suspended POC accounted for roughly 1% of

Why Dissolved Organics Matter

DOC, |LiM 0

20

40 I

60 ^1

80

100

Figure 2 Vertical profile of DOC in the southwest Atlantic Ocean (after Menzel and Ryther, 1968). Arrow lengths indicate the range of measured values, with the profile line passing through the mean value for that depth. Values in parentheses represent the total number of multiple analyses at one depth.

DOC below 500 m depth and also was essentially invariant. Dissolved O2 varied linearly versus salinity (Fig. 3) at the core of Antarctic intermediate water in all profiles. This observation of minimal DOC variation (Fig. 2) over a substantial oxygen gradient of > 100 /xM (Fig. 3) supported previous evidence for Httle or no DOC respiration below 500 m (Menzel and Ryther, 1970). By comparison, the theoretical OC/O2 ratio for respiration of "average marine plankton" is 106/138 = 0.77 (Redfield et al, 1963), whereas the best current estimate is near 0.70 (Anderson, 1995). Russian researchers at this time were measuring DOC concentrations by high-temperature combustion of freeze-dried samples. Although this method indicated concentrations that were approximately three times higher than those obtained with persulfate (see review by Starikova, 1970), minimal changes in deep-ocean DOC profiles were nonetheless noted (Skopintsev, 1966). Independent evidence that deep-sea DOM is refractory came from a variety of other sources. Barber (1968) demonstrated that DOM concentrated fivefold from deep-ocean water was not measurably utilized by marine bacteria and argued

John I. Hedges JUU

-^

F—"^^T—•^—T

H

-1

250 m''-,.

"X,

200

•' 1

150 0.200

^ ^ ^ ^ ^ 1

1

1

^^^^^^^^^

0.300 0.400 Salinity - 34.000

0.500

Figure 3 Measured dissolved oxygen versus the salinity in excess of 34.000 at the core of the Antarctic intermediate water in the southwest Atlantic Ocean (based on data from Table 1 of Menzel and Ryther, 1968). The equation of the best-fit line (r^ = 0.992) to the 13 data points is O2 = (-439 ± 12) x (S-34) + (374 ± 4).

against previous speculation (Jaanasch, 1967) that seawater DOM might simply be too dilute to serve as a suitable substrate. P. M. Williams (1968b) showed that the stable carbon isotopic composition of DOC is consistent with a predominantly marine origin and essentially constant throughout the water column of the San Diego trough. The definitive experiment of the decade, however, was the demonstration by Williams et al (1969) that the radiocarbon content of dissolved organic matter from the deep Pacific Ocean corresponds to a radiocarbon "age" of roughly 3400 years BR If this radiocarbon "age" is assumed to represent a mean residence time (Williams et al, 1969), it corresponds to a steady-state flux of roughly 0.2 x 10^^ g C/year through the ocean DOC pool (650-700 x 10^^ g C). Critically, this small flux would necessitate that only 0.4% of global primary production enters the marine DOC pool per year. Although a flux of this order could be supported by riverine DOC discharge alone (Williams, 1971; Mantoura and Woodward, 1983), the stable carbon isotopic composition of seawater DOC points toward a marine origin (WilHams, 1968a). WiUiams (1971) concluded that the predominant uncharacterized fraction of seawater DOM is humic-like and thus intrinsically unreactive. At the same time, parallel evidence was accumulating that an appreciable fraction of DOM in surface ocean waters can be physically and biologically reactive under at least some conditions. Natural slicks were observed to form and disperse

Why Dissolved Organics Matter rapidly at the ocean surface (Ewing, 1950; Jarvis, 1967) and to contain a variety of surface-active organic materials (Garrett, 1967, 1970) that could be concentrated by a dipped screen (Garrett, 1965), rotating drum (Harvey, 1966) or, "bubble microtome" (Maclntyre, 1966). In a series of experiments, Sieburth and Jensen (1968, 1969) demonstrated exudation of DOM by phaeophyta (kelp) and associated formation of sea surface slicks (Sieburth and Conover, 1965). Duursma (1961,1963, 1965) observed greater than twofold seasonal variation of DOC in surface waters of the North Sea. This indication of cycling on a monthly time scale suggested the possible use of DOC as an indicator of primary production. In contrast to results for deep water, Barber (1968) found that DOM concentrated from surface seawater exhibited a relatively short half-life (1-2 months) with respect to bacterial remineralization. The list of chromatographically measured biochemicals also increased substantially and the more abundant fatty acids, amino acids, and sugars had been quantified in surface waters and over a few deep-sea profiles (Holm-Hansen et al, 1966; Duursma, 1965; WiUiams, 1971). However, only about 10% of the DOC in surface and subsurface waters could be accounted for as individually measurable biochemical types, even when results from separate studies were added together (Williams, 1971). Although potentially labile biochemicals were evident, their low concentration was taken as additional evidence for a largely refractory pool of bulk DOM. The decade closed with a short conmiunication by Riley and Taylor (1969) describing how fatty acids and humic substances can be recovered from acidified seawater (pH 2) by sorption onto a cross-linked polystyrene resin called Amberlite XAD-1.

III. DOM RESEARCH IN THE 1970s The perception of a labile DOM component in the surface ocean accompanied by largely inert DOM (marine humus) that predominates below ~500 m continued to develop in the 1970s. The decade opened with the report by Williams and Gordon (1970) that the stable isotope composition of DOC at multiple stations in the northeast Pacific Ocean is remarkably uniform (5^^C = -22.6 db 0.6%^ ) and independent of depth and time, as well as dissolved O2 and DOC concentrations. The observation that these values were similar to those of local POC and marine plankton pointed toward a predominant marine origin of oceanic DOM. This inference was supported by a very different 8^^C value of —28.5%o measured for DOM from the Amazon River (Williams, 1968a). Although rivers discharge DOC at a rate sufficient to support the entire marine pool (WiUiams, 1971), the much more ^^C-enriched composition of marine DOC indicates that land-derived DOC must be rapidly removed or profoundly changed in its stable carbon isotopic composition. Minimal changes in the S^^C of marine DOC in depth profiles

7

John I. Hedges Ox

lUUI

80

2

=1

d o Q

* 0

• °\

•

•

•

40

A 24.0-24.5 I 1 • 25.0 - 25.5 1 D 25.5-26.0 1 • 26.0 - 26.5 1 A 26.5 - 27.0 . „ , ^ ^^ ^ 1 X >27.5 -^^^O-^^-^l

3
23

Baltic Sea Bothnian Sea 63" 30'N. 19" 48'E Gulf of Riga 58" 15". 24" 24' E Gulf of Riga 57" 16'N, 24" 23' E

4

ADOC

6G110

3-20

ADOC

70

3-20

ADOC

630

&loo

ADOC

Reference Daly et al., 1999

Daly et al., 1999

Range of maximum DOC from a 2-year time series; ADOC, change in DOC between spring and summer ADOC, change in DOC between spring and summer transect means; coastal input exists ADOC change in DOC between spring and summer transect means; coastal input exists

Zweifel et al., 1993

1-Year time series; ADOC, change in DOC between February and November; range, maximum concentration over 100 m

Copin-MontCgut and Avril, 1993

Zweifel. 1999

Zweifel, 1999

Mediteranean Sea 43" 25' N, 07" 52' E

1233

8-2 1

North Atlantic English Channel 48" 45' N, 3" 57'W

0-30

275

70

ADOC, change in DOC between April and August

Wafar et al., 1984

ADON

3

Wafar et al., 1984

ADOP

0.12

ADON, change in DON between August and September ADOP, change in DOP between August and October 1-Year time series; ADOC, change in DOC between March and June

ADOC

2100'

English Channel 50" 02' N, 4" 22' w

0-70

125

ADOC

1820b

26

English Channel 50" 02' N, 4" 22' w Norwegian Sea 66" N, 2' E

0-70

125

ADON

119'

1.7

0-50

ADOC

165O-418Ob

33-83

5

ADOC

108

45

ADOC

15

ADOC

North Sea 53" Ol'N, 4" 21'E North Sea 58" 55' N, 0" 32' E Bedford Basin, Nova Scotia, Canada

ADOC, change in DON between March and June

Wafar et al., 1984 Banouh and Williams, 1973; PP from Boalch et aL, 1978 Banoub and Williams, 1973

Annual range at for a 3-year time series; no systematic variability with DON 1-Year time series; ADOC, change in DOC between March and April

Bersheim and Myklestad, 1997 Duursma, 1963

21

Measured as dissolved carbohydrates

Ittekkot et al., 1981

40

ADOC, change in DOC between February and April; Coastal hay

Kepkay et al., 1997 (Continues)

Table I (Continued) ADOM stock and concentration

DOM depth'(m)

PP (mmol m-2 day-')

ADOM type

mmolm-2

uM

Sargasso Sea 31" 50'N, 64" 10'W

G250

57-123'

ADOC

500-1400

2 4

Range of maximum DOC from an annual range at for a 5-year time series; no systematic variability with DON

Strait of Georgia

G20

100

ADOC

2500

125

Strait of Georgia

c-20

100

ADON

I-Year time series, ADOC, change between February and August Calculated from mass balance considerations

Site

Southern Ocean Antarctic Polar Front Zone (APFZ) Australian sector 56'45" 24' S Australian sector 56"45O 24' S Atlantic sector 48-52O.5, 2640" w

13

Comments

Reference Carlson et aL, 1994; Hansel1 and Carlson, 2000 Parsons et al., 1970 Williams, 1995; ADON estimates

Surface layer

ADOC

5-15d

Calculated as increase above deep water concentrations;range represents spatial variability

Ogawa et al., 1999

Surface layer

ADON

1.5-7.2d

Calculated as increase above deep water concentrations;range represents spatial variability

Ogawa et aZ., 1999

Calculated from difference max, and minimum concentration in APFZ, Largest accumulationnear southern periphery of A P E

Dafner, 1992

G50

66

ADOC

5ood

Atlantic sector 47-60"s

0-100

Indian Ocean sector 49-63" S

0-100

Antarctic Continental Shelf Systems Bransfield Strait off Palmer Peninsula Prytz Bay 68" 30'S, 77" 50' E Ross Sea 76" 30' S transect line Ross Sea 76" 30's transect line

6-2 1

ADOC

< 1-2od

ADOC

4-16d

0-100

ADOC

>500-1000

15

ADOC

>go00

ADOC

370-1 140

0-150

Surface

80-226e

ADON

0-23

0.1-5.5

Calculated as increase above deep water concentrations; range represents spatial variability Calculated as increase above deep water concentrations; range represents spatial variability

K&kr et al., 1997 Wiebinga and de Baar, 1998 & citations within

Bloom of Phaeocystis sp., Thalassiosira sp., and Corethron sp.

Bolter and Dawson, 1982

Phaeocystis antarctica bloom

Davidson and Marchant, 1992 Carlson et al., 2000

Time series measurement of a composite growing season; ADOM, change between Oct. and Jan. transect means for surface 150 m ADON represent change in surface 150 m; mean C:N ratio of ADOM = 6.2

Carlson et al., 2000

Note. PI' was integrated over the euphotic zone of each site. Blank space means data not available. Table is expanded from Carlson et al. (2000). aDOM depth refers to depth where sample was collected or depth used to integrate DOM stock. bCalculated from integration depth and mean ADOM concentration for given depth horizon. 'Primary production integrated over 140 m. 'Winter and early spring DOM concentrations in Southern Ocean equal deep DOM concentrations (Kahler et al., 1997; Carlson et al., 2000) due to deep mixing and remineralization; thus, during the growing season DOM concentrations in excess of deep water values are assumed to be seasonally produced. 'Primary production estimates from Smith and Gordon (1997) and Smith et al. (2000).

98

Craig A. Carlson

A. EXTRACELLULAR PHYTOPLANKTON PRODUCTION

Over four decades ago, extracellular release of carbohydrates was identified in algal cultures (Lewin, 1956; Guillard and Wangersky, 1958). Since then there has been an explosion of research on extracellular phytoplankton production of carbohydrates, nitrogenous compounds, and organic acids. Several extensive reviews are now available that discuss the rates and potential physiological mechanisms of algal release of DOM in marine systems (Fogg, 1983; Williams, 1990; Baines and Pace, 1991; Nagata, 2000). This topic will be discussed briefly here. Direct measurements of bulk DOM or specific compounds as well as radioisotope techniques have been employed to study ER. Tracing the uptake of ^"^C bicarbonate by phytoplankton and release into DO^'^C is often used as a method for assessing extracellular C production (Fogg, 1966). Methodologically this technique is easy; however, many artifacts associated with this method can bias the interpretation of the data. For example, a lag in DO^^C release can occur because intracellular pools of organic metabolites do not immediately reach isotopic equilibrium; thus, if DOC labeling rates are calculated with a constant tracer release model then actual release rates will be underestimated (Lancelot, 1979; Smith, 1982). In addition, uptake of DO^'^C by heterotrophic microorganisms can result in a decrease in measured DO^'^C release over an incubation (Wiebe and Smith, 1977; Lancelot, 1979), leading to an underestimate of actual DO^'^C release. Alternatively, overloading cells on afilter,rupturing cells during filtration and mishandling of sample can lead to overestimates of ER, especially in oligotrophic systems (Sharp, 1977; Goldman and Dennett, 1985). Finally, the appearance of DOM in incubated seawater samples is difficult to attribute solely to extracellular release due to the presence of a mixed microbial assemblage present and the potential contribution of other DOM production processes within the incubation bottles. Nonetheless, theory (Bj0msen, 1988) as well as field and experimental evidence (Table 11; Mague et ai, 1980; Fogg, 1983; WiUiams, 1990; Baines and Pace, 1991; Karl et al, 1998) suggests that ER of DOC is a normal function of healthy in situ photoautotrophic growth. Extracellular release of dissolved organic nitrogen (DON) has also been assessed, using ^^N tracer techniques (Bronk and Gilbert, 1993; Hu and Smith, 1998; Bronk and Ward, 2000), and is addressed in this book (see Bronk, Chapter 5). Percent extracellular release (PER) is a measure of the ^"^C accumulating in DOC relative to total particulate plus dissolved primary production following incubation. In a review of culture experiments, Nagata (2000) reported that during exponential growth PER values averaged 5% (typical range 2-10%) for a variety of marine phytoplankton isolates. Sharp (1977) criticized PER as a useful indicator in the field, suggesting that procedural artifacts such as inadequate assessment of control blanks and rupturing of cell during processing could lead to artificially high

Production and Removal Processes

99

PER, especially in systems where PP is low. The body of hterature regarding DO^'^C release in nature is large, growing, and often conflicting as to its importance (Sharp, 1977; Fogg, 1983;Bj0msen, 1988; Williams, 1990). Table II demonstrates the wide range in ER rates (0-12 jig C L-^h-^) and PER (0-80%) observed in the field for a variety of coastal and oceanic systems. Several factors, such as community structure, light intensity, nutrient deficiency, and temperature, affect PER in situ. However, for any given controlling factor one can find examples of contrasting effects on PER (Table III), indicating the complex interactions of phytoplankton C production and environmental conditions. Although physiological state, species composition, and local chemical/physical conditions can significantly affect PER, there is little systematic variability of PER across productivity regimes (Baines and Pace, 1991; Nagata, 2000). In a cross-system analysis, Baines and Pace (1991) found that, while the absolute rate of ER varied depending on the nutrient regime of the system, the average PER was 13%. 1. Extracellular Release Models Two models have been proposed to explain extracellular production by photoautotrophs. They are the overflow model (Fogg, 1966, 1983; Wilhams, 1990; Nagata, 2000) and iht passive diffusion model (Fogg, 1966; Bj0msen, 1988). a. Overflow Model Fogg (1966) reasoned that because photosynthesis is largely regulated by irradiance and cellular growth is constrained by the availabihty of inorganic nutrients, a cell's photosynthate may be produced faster than it is incorporated and would therefore be actively released via an "overflow" mechanism. It may be energetically less costly for a cell to discard surplus nonnitrogenous compounds (i.e., capsular material and carbohydrates) than to store it under nutrient-limiting conditions (Wangersky, 1978; Wood and Valen, 1990). According to the overflow model, DOM exudation should (1) correlate to the photosynthetic rate, (2) be absent at night, and (3) be composed of both low-molecular-weight (LMW < 1000 Da) and high-molecular-weight (HMW> 1000 Da) DOM (Fogg, 1966; Bj0msen, 1988; Williams, 1990). Factors such as light intensity and nutrient availabihty potentially control the degree at which the overflow model functions. However, Bratbak and Thingstad (1985) used a modeling exercise to point to a paradoxical situation in which nutrient-stressed phytoplankton appeared to stimulate bacterial production and thus increased competition for nutrients. Bj0msen (1988) argued that active release of extracellular DOC would exacerbate a nutrient limiting scenario and suggested that the release of DOM from phytoplankton was a passive process.

•a •c o

I o G

o

s 2 2 OS

PC

w CU

(U

's

2

-l

• S- ,

Q

^

K

l>5

§ g

P 6 ^ 13 S ^ ^

Q Q Q

u u u o oo

^^ s ^ ^ i:^ U O

O

^II "K

u 1 :z; ^^ PQ U

'-' i i

3

3

l>3

(/I

II

C«

d d

u :^ o o

d

3

T3

^

Q Q Q

u u :z; o oo

» H

3

«3 O

c o

^

o N

§ _c &

3 *« O

^

§3

>

o c

1^ 5

c

G

>

10 years (Jenkins, 1980) indicating that semilabile DOM advected along isopycnal surfaces to these depths would have turnover rates on the time scale of many years. The continuum of reactivity of the semilabile pool also varies between ocean systems (Carlson et al, 2000). The relative contribution of specific compounds,

131

Production and Removal Processes

iii

• •

Ross Sea

Sargasso Sea

Semi-labile < lyear Semi-labile > lyr Refractory

Mediterranean Sea

Figure 7 Contribution of semilabile DOC to the bulk DOC in the surface 250 m of the Ross Sea and the Sargasso Sea and surface 100 m of the Mediterranean Sea. All stocks were integrated vertically then normalized to integration depth for comparison purposes. Black shaded area of each column represents refractory DOC. DOC concentrations below 1000 m at each study site were used to represent refractory DOC contribution. The sum of the gray areas of each column represents semilabile DOC (i.e., the integrated DOC stock in excess of refractory DOC stocks). The light gray represents the proportion of semilabile DOC that turns over on time scales within 1 year as determined from the difference between the annual maximum and minimum stocks within the depth horizon of each time-series study site. The dark gray area represents the portion of semilabile DOC that is in excess of refractory DOC but does not vary over the time scale of 1 year (i.e., turns over on time scales of > 1 year). Estimates of turnover are based on observed changes in integrated pools from three time series studies. The Ross Sea data were adapted from Carlson et al. (2000); the Sargasso Sea entry was determined from the 1995 spring phytoplankton bloom event (Hansell and Carlson 1998b); the Mediterranean Sea example was adapted from Copin-Montegut and Avril (1993).

such as aldoses, can be used as an index of DOM reactivity (Skoog and Benner, 1997; Biersmith and Benner, 1998). Tlie semilabile DOC pool observed in the Ross Sea, Antarctica, contains a higher percentage of aldoses (up to 50%) than the Arctic, equatorial Pacific, and the Sargasso Sea (Fig. 8; Kirchman et al, 2001). The large contribution of reactive compounds in Ross Sea DOM is consistent with rapid turnover of semilabile DOM observed there (Carlson et al, 2000; Kirchman et al, 2001). Carlson et al (2000) found that nearly the entire semilabile DOC pool present in the Ross Sea in February 1997 (1.14 mol C m"^) was consumed within 2 months. In contrast, only 70 and 20% of the total semilabile DOC pool turned over on the time scales of less than 1 year for the Mediterranean Sea (CopinMontegut and Avril, 1993) and the Sargasso Sea (Carlson et al, 1994; Hansell and Carlson, 2001), respectively (Fig. 7). The inorganic nutrient regime may be an important factor in controlling the stoichiometry of freshly produced DOM and, in turn, its lability (Williams, 1995;

Craig A. Carlson

132 120 c3

u o Q

100 H

80-4

-]

60 H

-^ •S

40J

o

20-

T

1 _L T

jy -j

Arctic

1

^ ^ ^1

, Ross Sea 1 EqPacSl , EqPac S2 1Sargasso Sea1

Ocean Site Figure 8 Dissolved combined aldoses as a fraction of semilabile DOC in surface waters of various oceanic systems. Data from the Equatorial Pacific are from D. Kirchman and N. Borch (unpublished data), the central Arctic from Rich et al. (1997), the Ross Sea from Kirchman et al. (2001), and the Sargasso Sea from C. Carlson and M. Otero (unpublished data). The top, bottom, and line through the middle of each box represent the 75th, 25th, and 50th percentiles, respectively. The lines on the top and bottom of each box extend from the 10th to the 90th percentile of the data. Figure adapted with permission form Kirchman et al. (2001).

Ogawa et al, 1999; Carlson et al, 2000). Carbon-rich DOM (C:N >12 for semilabile DOM) in the North Atlantic was resistant to microbial degradation on seasonal time scales (Williams, 1995; Hansell and Carlson 2001; Kahler and Koeve, 2001), whereas nitrogen-rich DOM (C:N of ^6.7 for semilabile DOM) production in the nutrient replete Ross Sea was utilized on time scale of weeks (Carlson et al, 2000). The variable composition of semilabile DOC within and between ocean sites indicates that applying a single decay constant to calculate turnover of the integrated semilabile DOC pool is inappropriate in most cases. Semilabile DOM turnover is often calculated from integrated DOM stocks and instantaneous BP rates (i.e., semilabile DOM turnover = semilabile stock/(BP/BGE). However, this method of calculating turnover rates probably overestimates the actual rates because instantaneous BP rates (determined from [^H] thymidine or [-^H] leucine incubations) probably do not reflect growth supported by recalcitrant material. Instantaneous BP rates are more likely to be an index of growth supported by the rapid flux of labile DOM rather than semilabile DOM. A^'^C evidence does indicate that

Production and Removal Processes

133

bacterioplankton are able to take up "old" DOM (Cherrier et ah, 1999); however, remineralization of semilabile DOM can be slow and at times undetectable on time scales of days to weeks (Fig. 6b; Carlson and Ducklow, 1996; Cherrier et al, 1996; Carlson et ah, submitted for publication). Thus, one should not assume that semilabile DOM is utilized at a rate comparable to instantaneous BP measurements.

V. DOM ACCUMULATION Why does DOM accumulate? In the open ocean the net production of DOC is ultimately due to the decoupling of biological production and consumption processes. While there are several DOM production mechanisms (see section II), the dominant oxidizers of marine DOM are heterotrophic bacterioplankton (Azam and Hodson, 1977). Thus, factors that prevent rapid microbial utilization of "freshly produced" DOM result in its accumulation. These factors may include: (A) abiotic transformation of labile components to biologically recalcitrant compounds, (B) biological production of recalcitrant DOM, and (C) limitations on heterotrophic bacterial growth.

A. ABIOTIC FORMATION OF BIOLOGICALLY RECALCITRANT DOM Refractory or recalcitrant DOM may be formed from labile compounds by either cross-linking polymerization of LMW DOM (condensation reaction catalyzed by light and metals; Harvey et al, 1983) or modification of LMW labile material (e.g., proteins Hedges, 1988). Keil and Kirchman (1994) demonstrated that labile organic matter could be modified abiotically to a form resistant to rapid microbial oxidation. Condensation reactions, binding of monomers to macromolecular DOM (Carlson et al, 1985), adsorption to colloids (Kirchman et al, 1989; Keil and Kirchman, 1994; Nagata and Kirchman, 1996), and exposure to UV irradiation (Keil and Kirchman, 1994;Naganuma^fa/., \996\ Gohl^v et al, 1997; Benner and Biddanda, 1998; Tranvik and Kokalj, 1998) have all been proposed as mechanisms that physically alter DOM to a molecular structure that impedes DOM degradation. The effects of UV exposure on DOM are complex and seemingly yield both labile (Kieber et al, 1989; Moran and Zepp, 1997) and recalcitrant DOM products (Keil and Kirchman, 1994;Naganuma^ffl/., 1996;Gobler^/(3/., 1997; Benner and Biddanda, 1998; Tranvik and Kokalj, 1998; see Mopper and Keiber, Chapter 9). Benner and Biddanda (1998) found that exposure of euphotic zone DOM to UV irradiation reduced bacterial production by 75% while exposure of deep DOM (150-1000 m) to UV enhanced bacterial production by 40%. They concluded that the chemical composition of DOM dictates whether phototransformations produce

134

Craig A. Carlson

bioavailable or bioresistant compounds. The exact mechanisms and magnitude of these abiotic transformations remain unknown. Studying abiotic transformation of DOM from labile to recalcitrant forms may provide clues as to the origin of refractory DOM.

B. BiOTic FORMATION OF RECALCITRANT D O M Abiotic processes lead to the restructuring of recognizable labile compounds into complex macromolecules that are not recognized by traditional chemical analysis. These processes appear to "shield" the labile component from biological oxidation (Keil and Kirchman, 1994). Recent studies have also identified unmodified recalcitrant components of DOM, formed by direct biosynthesis, that can contribute a large fraction of the HMW DOM found in the surface waters (Tanoue et ai, 1995; Aluwihare et al, 1997; McCarthy et ai, 1998). Eukaryotic and prokaryotic organisms are both potential sources of biologically recalcitrant DOM. 1. Eukaryotic Sources Extracellular release is a major source of carbon rich carbohydrates in the surface ocean. Polysaccharides contribute a major fraction of the HMW (> 1000 Da) DOM (Benner et al, 1992; McCarthy et ai, 1996; Skoog and Benner, 1997). In a culture experiment with marine diatoms, Lara and Thomas (1995) observed an increase in recalcitrant DOC as POC decreased, indicating that cellular components such as cell wall material may be the source. A compound resembling acyl heteropolysaccharide (APS), a metabolically resistant and dominant polysaccharide in the surface ocean (Aluwihare etal, 1997), has been shown to be released directly by marine diatoms and haptophytes (Aluwihare and Repeta, 1999). This APS-like polysaccharide had a slower degradation rate relative to the total polysaccharide fraction of the phytoplankton exudate. Biersmith and Benner (1998) found similarities between the aldose signature of HMW phytoplankton exudate and HMW DOM isolated from various locations in the surface ocean. Phytoplankton community structure may play an important role in the production of recalcitrant DOM. Aluwihare and Repeta (1999) found that in three species of phytoplankton studied, all produced APS-like polysaccharides. However, the percentage of polysaccharides released as APS varied considerably between species. 2. Prokaryotic Sources Prokaryotes can also be sources of recalcitrant DOM (Table VI). Brophy and Carlson (1989) observed bacterial transformation of ^"^C labeled glucose

Production and Removal Processes

135

and leucine and subsequent release of recalcitrant HMW (700-1400 Da). Similarly, Tranvik (1993) found that approximately 3% of the initial glucose concentration was transformed into humic-like DOM after a 1-week incubation. Heissenberger et al. (1996) reported ^^C-labeled leucine was transformed into recalcitrant HMW (> 50,000 Da) DOM in a bacterial batch culture. A subsequent study hypothesized that the HMW material produced by microbial growth was capsular material (Stoderegger and Hemdl, 1998). Ogawa et al. (2001) found that when glucose and glutamate were added to bacterial cultures the labile compounds were utilized rapidly; however, the bacterioplankton also produced DOM that resisted further microbial degradation for time scales of more than 1 year. These studies attributed recalcitrant DOM production to bacterial processes, but they were not able to rule out the possibility that viral lysis or grazing contributed to the HMW DOM production. Nonetheless, DOM with bacterial-like biochemical characteristics is ubiquitous in the surface ocean, indicating a bacterial source for some recalcitrant DOM (Tanoue et al, 1995,1996; McCarthy et al, 1998). Tanoue et al. (1995) identified a dissolved protein, homologous to the Gramnegative bacterial membrane porin P, as being common to various ocean basins. Porins are resistant to proteases and rapid microbial degradation (Tanoue et al, 1996 and citations within). Peptidoglycans, the main structural component of bacterial cell walls, have also been proposed as the likely source of enriched D-enantiomer amino acids (D-amino acids) found in two oligotrophic sites (McCarthy et al, 1998). The polysaccharide matrix, with its unusual peptide structures, yields a polymer that is resistant to many common hydrolytic enzymes, rendering it bioresistant (McCarthy et al, 1998). Liposome-Uke particles (aqueous compartments enclosed by a lipid bilayer) released from bacterioplankton via viral lysis or nanoflagellate grazing may be an additional prokaryotic source of recalcitrant DOM (Nagata and Kirchman, 1992; Nagata, 2000). The exact mechanism by which these bacteria-associated compounds enter into the dissolved phase (release from bacteria directly or a byproduct of microzooplankton grazing or viral lysis) are not well understood or quantified. Nonetheless, these bacteria-derived compounds are now becoming recognized as an important source of recalcitrant or refractory DOM. C. LIMITATION OF BACTERIAL GROWTH AND ACCUMULATION OF BIODEGRADABLE

DOM

The accumulation of DOM in surface seawater has been largely attributed to the accumulation of recalcitrant material resistant to microbial degradation (Billen and Fontigny, 1987; Brophy and Carlson, 1989; Legendre and Le Fevre, 1995; Tanoue et al, 1995; Carlson et al, 1998). Based on the assumption that accumulated

136

Craig A. Carlson

DOM is recalcitrant, one might expect that growth of heterotrophic bacterioplankton is initially limited by the availability of labile DOM (Kirchman et al, 1990; Kirchman, 1990; Keil and Kirchman, 1991; Carlson and Ducklow, 1996; Cherrier et al, 1996; Carlson et al, in press; WilUams, 2000). However, the view of labile DOM limitation of bacterial growth has recently been challenged by the "malfunctioning microbial loop" hypothesis (Thingstad et al, 1997). This hypothesis states that competition for limiting nutrients (Bratbak and Thingstad, 1985; Zweifel et al, 1993; Thingstad and Rassoulzadegan, 1995; Cotner et al, 1997; Rivkin and Anderson, 1997; Thingstad et al, 1998) and grazing pressure (Thingstad and Lignell, 1997; Zweifel, 1999) reduce bacterioplankton growth rate, biomass, and carbon demand to levels that allow accumulation of biodegradable DOC during biologically productive seasons. Low temperatures have also been suggested as a mechanism that inhibits BP (Pomeroy and Deibel, 1986; Pomeroy et al, 1991; Shiah and Ducklow, 1994) and may foster DOM accumulation (Zweifel, 1999). However, Carlson et al (1998) and Ducklow et al (in press), found little evidence to support the hypothesis of temperature regulation on bacterial growth or DOC accumulation in the Ross Sea, Antarctica. However, temperature regulation may be more important in systems that demonstrate large seasonal temperature ranges. According to the "malfunctioning microbial loop" hypothesis, one would expect that by reducing grazing pressure and adding potentially limiting nutrients to dilution cultures, bacterioplankton growth and DOC utilization would be enhanced. The experimental results of Zweifel and Hagstrom (1995), conducted in the Baltic Sea, support this hypothesis by showing enhanced bacterial growth and DOC utilization in cultures amended with inorganic N and P. In contrast to these findings, inorganic amendments had neither an effect on bacterial production nor DOC remineralization in the oceanic eastern North Pacific (Cherrier et al, 1996) or the northwestern Sargasso Sea (Carlson and Ducklow, 1996; Carlson et al, in press). S0ndergaard et al (2000) found that inorganic nutrient amendment had only a marginal effect on DOC degradation. In his review on the controls of microbial growth, Williams (2000) found that the nutrient limitation hypothesis appeared to be more frequently sustained in coastal regions (see his Table IV) and to a lesser extend in oceanic waters. This is not to say that inorganic nutrient limitation of bacterioplankton growth does not occur in oceanic waters. In fact, several studies in the northwestern Sargasso Sea have demonstrated that, at times, bacterioplankton production can respond to amendments of inorganic nutrients (Cotner et al, 1991 \ Rivkin and Anderson, 1997; Caron et al, 2000). However, in studies where DOC consumption was measured directly, no evidence exists to suggest that amending surface water assemblages with inorganic macronutrients further reduces semilabile DOC concentration below the mean mixed layer concentrations of the northwestern Sargasso Sea on the time scales of weeks (Carlson et al, in press).

137

Production and Removal Processes D. MICROBIAL COMMUNITY STRUCTURE AND DOM

UTILIZATION

In addition to the inorganic nutrient regime, the structure of the microbial community may play an important role in regulating the accumulation and subsequent remineralization of semilabile DOM. For example, in the northwestern Sargasso Sea DOC stocks accumulate rapidly within the euphotic zone shortly after water column stratification and persist at elevated concentrations throughout the summer into early autumn (Carlson et ai, 1994; Hansell and Carlson, 2001). During seasonal overturn a portion of the seasonally accumulated semilabile DOC can be exported to depths >200 m. Once isolated within the aphotic zone the exported DOC is remineralized relatively quickly on time scales of weeks to months (Carlson et ai, 1994; Hansell and Carlson, 2001). Why does the seasonally produced semilabile DOM escape rapid microbial degradation in the surface but become available to microbial remineralization at depth? While inorganic macronutrients are found at elevated concentrations at depth, simply amending surface water microbial assemblages with inorganic macronutrients did not appear to stimulate DOC removal in experimental cultures conducted in the northwestern Sargasso Sea (Carlson et al, in press). Prokaryotic phylogenetic diversity is greater below the euphotic zone compared to the surface waters (Giovannoni etal, 1996; Gordon and Giovannoni, 1996). Archaea dominate in the mesopelagic regions of some ocean sites (Kamer et al, 2001). Kamer et al. (2001) proposed that the high percentage of Archaea cells containing significant amounts of rRNA suggests that they are metabolically active. Do some Archaea specialize in utilizing diagenetically altered semilabile DOM? Vertical gradients in the availability of nutrients and energy may be responsible for the observed diversification and specialization of microbial communities. These specialized microbial communities may regulate consumption of semilabile DOC transported to depth. Further experimental work is necessary to gain insight and to quantify potential linkages between specialized microbial assemblages and biogeochemical processes, such as utilization of semilabile DOC.

VI. SUMMARY In this chapter, I have outlined our present understanding of the DOM production and removal processes, the characteristics of the general pools of lability of the bulk DOM pool, and factors that lead to DOM accumulation. 1. DOM production mechanisms include direct phytoplankton release, zooplankton-associated processes (i.e., grazing and excretion), virus and bacteriainduced release, solubiHzation of particles, and prokaryotic DOM production.

138

Craig A. Carlson

Experimental and field evidence suggests that extracellular release of DOM is a normal function of phytoplankton production with typically 13% of PP being released as DOM; however, significant variability exists in the literature (Table II). The magnitude of extracellular release is dependent on a variety of physiological and environmental conditions not fully understood yet. Empiricists (Jumars et al, 1989; Nagata, 2000) and modelers (Anderson and Ducklow, submitted for publication) both suggest that under steady-state conditions the bulk of DOC supply comes from zooplankton processes. Nagata (2000) suggests that as much as 30% of PP is released from protozoan herbivory with an additional contribution from macrozooplankton grazing and bacterivory. Significant study is still required to properly assess the contribution of DOM production via viral impact, solubilization of particles and direct release of organics from prokaryotes. 2. Bacterioplankton (or prokaryotic) oxidation of DOM is considered the main sink for recently produced DOM; however, the role of UV oxidation is now recognized as an important removal process especially for refractory DOM. Sorption of DOM onto sinking particles is also recognized as a potential DOM removal mechanism within the oceans interior. Work continues toward trying to identify and quantify processes that remove refractory DOM. 3. The bulk DOM pool represents a broad continuum of biological lability ranging from material that turns over on time scales of minutes to days (labile DOM), to material that turns over on time scales of weeks to years (semilabile), to material that survives for decades to millennia (refractory DOM). The refractory pool represents the majority of DOC present in the surface waters of thermally stratified waters (^70% in temperate and tropical waters). The labile pool is kept at low concentrations due to high turnover by microbial activity. The majority of the vertical structure in a DOM profile in thermally stratified systems is composed of semilabile DOM, which accumulates in excess of the refractory background DOC stocks. Factors such as biological conmiunity structure and nutrient regime may play a role in the production of semilabile DOM. 4. Accumulation of DOC results from the uncoupling of DOM production and removal processes. The production of biologically resistant compounds via physical processes such as condensation reactions or phototransformation can result in the production of biologically resistant DOM. Unmodified recalcitrant components of DOM, formed by direct biosynthesis, has also been identified for both phytoplankton and bacterioplankton. Inorganic nutrient limitationcontrol of DOM accumulation remains an interesting hypothesis but evidence of its support is ambiguous. Finally spatial variability of microbial conmiunity structure may also play a role in the processing and cycling of recalcitrant compounds.

Production and Removal Processes

139

ACKNOWLEDGMENTS I particularly express my gratitude to Hugh Ducklow and Dennis Hansell, who have been great collaborators and friends along this path. This chapter benefited greatly from reviews and discussions by and with James Christian, Hugh Ducklow, David Smith, David Kirchman, and Dennis Hansell. Thanks to Walker Smith, David Smith, and Deborah Steinberg for access to unpublished data. I thank Stuart Goldberg and Rachel Parsons for assistance in generating some of the table data used in this chapter. This work has been supported by NSF Grants OCE 9617795, OCE 9619222, MCB-9977918 and OCE-0196305. This is U.S. JGOFS contribution 712.

REFERENCES Aluwihare, L. I., and Repeta, D. J. (1999). A comparison of the chemical characteristics of oceanic DOM and extracellular DOM produced by marine algae. Mar. Ecol. Prog. Sen 186, 105117. Aluwihare, L. I., Repeta, D. J., and Chen, R. F. (1997). A major biopolymeric component to dissolved organic carbon in surface sea water. Nature 387,166-169. Amon, R. M. W., and Benner, R. (1994). Rapid cycling of high-molecular-weight dissolved organic matter in the ocean. Nature 369,549-552. Amon, R. M. W., and Benner, R. (1996). Bacterial utiUzation of different size classes of dissolved organic matter. Limnol. Oceanogn 41,41-51. Andersen, O. K., Goldman, J. C , Caron, D. A., and Dennett, M. R. (1986). Nutrient cycling in a microflagellate food chain. III. Phosphorous dynamics. Mar. Ecol. Prog. Sen 31,47-55. Anderson, T. R., and Ducklow, H. W. (2001). Microbial loop carbon cycling in ocean environments studies using a simple steady state model. Aquat. Microb. Ecol. 26, 3 7 ^ 9 . Anderson, T. R., and Williams, P. J. 1. (1999). A one-dimensional model of dissolved organic carbon cycling in the water column incorporating combined biological-photochemical decomposition. Global Biogeochem. Cycles 13, 337-349. Andersson, A., Lee, C , Azam, R, and Hagstrom, A. (1985). Release of amino acids and inorganic nutrients by heterotrophic marine microflagellates. Mar. Ecol. Prog. Sen 23,99-106. Azam, F., and Cho B. C. (1987). Bacterial utilization of organic matter in the sea. In "Symposium of the Society for General Microbiology. 4.1. Ecology of Microbial Communities." University Press, Cambridge, England. Azam, R, Fenchel, T., Field, J. G., Gray, J. S., Meyer-Reil, L. A., and Thingstad, R (1983). The ecological role of water-column microbes in the sea. Mar Ecol. Prog. Sen 10, 257-263. Azam, R, and Hodson, R. E. (1977). Size distribution and activity of marine microheterotrophs. Limnol. OceanogK 22,492-501. Baines, S. B., and Pace, M. L. (1991). The production of dissolved organic matter by phytoplankton and its importance to bacteria: Patterns across marine and freshwater systems. Limnol. Oceanogn 36,1078-1090. Barber, R. T. (1968). Dissolved organic carbon from deep water resists microbial oxidation. Nature 220, 274-275. Banoub, M. W, and Williams, P. J. L. (1973). Seasonal changes in the organic forms of carbon, nitrogen and phosphorus in sea water at Ei in the English channel during 1968. /. Mar Biol. Assoc. UK 53,695-103. Bauer, J. E., Williams, P. M., and Druffel, E. R. M. (1992). ^"^C activity of dissolved organic carbon fractions in the north-central Pacific and Sargasso Sea. Nature 357,667-670.

140

Craig A. Carlson

Behrenfeld, M. J., and Falkowski, P. G. (1997). Photosynthetic rates derived from satellite-based chlorophyll concentration. Limnol Oceanogr. 42, 1-20. Benner, R. H. (2002). Composition and reactivity. In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. Hansell and C. A. Carlson, Eds.), pp. 59-90. Academic Press, San Diego. Benner, R., and Biddanda, B. (1998). Photochemical transformation of surface and deep marine dissolved organic matter: Effects on bacterial growth. Limnol. Oceanogr. 43,1373-1378. Benner, R., Pakulski, J. D., McCarthy, M., Hedges, J. I., and Hatcher, P G. (1992). Bulk chemical characteristics of dissolved organic matter in the ocean. Science 255,1561-1564. Bergh, 0., B0rsheim, K. Y., Bratbak, G., and Heldal, M. (1989). High abundance of viruses found in aquatic environments. Nature 340,467-468. Berman, T., and Holm-Hansen, O. (1974). Release of photoassimilated carbon as dissolved organic matter by marine phytoplankton. Mar. Biol. 28, 305-310. Berman, T., and Kaplan, B. (1984). Diffusion chamber studies of carbon flux from living algae to heterotrophic bacteria. Hydrobiologia 108,127-134. Biddanda, B., and Benner, R. (1997). Carbon, nitrogen and carbohydrate fluxes during the production of particulate and dissolved organic matter by marine phytoplankton. Limnol. Oceanogr. 43,506-518. Biddanda, B., Opsahl, S., and Benner, R. (1994). Plankton respiration and carbon flux through bacterioplankton on the Louisiana shelf. Limnol. Oceanogr. 39,1259-1275. Bidigare, R. R. (1983). Nitrogen excretion by marine zooplankton. In "Nitrogen in the Marine Environment" (E. J. Carpenter and D. G. Capone, Eds.), pp. 385-409. Academic Press, New York. Biersmith, A., and Benner, R. (1998). Carbohydrates in phytoplankton and freshly produced dissolved organic matter. Mar Chem. 63,133-144. Billen, G., and Fontigny, A. (1987). Dynamics of a Phaeocystis-dormndXQd spring bloom in Belgian coastal waters. II. Bacterioplankton dynamics. Mar Ecol. Prog. Ser 37, 249-257. Bj0msen, P. K. (1986). Bacterioplankton growth yield in continuous seawater cultures. Mar Ecol. Prog. Ser 30,191-196. Bj0msen, P. K. (1988). Phytoplankton exudation of organic matter: Why do healthy cells do it? Limnol. Oceanogr 33, 151-154. Bj0msen, P. K., and Kuparinen, J. (1991). Determination of bacterioplankton biomass, net production and growth efficiency in the Southern Ocean. Mar Ecol. Prog. Ser 71,185-194. Boalch, G. T., Harbour, D. S., and Butler, E. I. (1978). Seasonal phytoplankton production in the western Enghsh Channel 1964-1974. J. Mar Biol. Assoc. UK5H, 943-953. Bolter, M., and Dawson, R. (1982). Heterotrophic utiUzation of biochemical compounds in Antarctic waters. Neth J. Sea Res. 16, 315-332. Borch, N. H., and Kirchman, D. L. (1997). Concentration and composition of dissolved combined neutral sugars (polysaccharides) in seawater determined by HPLC-PAD. Mar Chem. 57, 85-95. B0rsheim, K. Y, and Myklestad, S. M. (1997). Dynamics of DOC in the Norwegian Sea inferred from monthly profiles collected during 3 years at 66°N, 2°E. Deep Sea Res. 744, 593-601. B0rsheim, K. Y, Myklestad, S. M., and Sneli, J. A. (1999). Monthly profiles of DOC, mono- and polysaccharides at two locations in the Trondheimsfjord (Norway) during two years. Mar Chem. 63, 255-272. Bratbak, G., Egge, J. K., and Heldal, M. (1992). Viral mortality of the marine alga Emiliana huxleyi (Haptophyceae) and termination of algal bloom. Mar Ecol. Prog. Ser 83,273-280. Bratbak, G., and Thingstad, T. F. (1985). Phytoplankton-bacteria interactions: An apparent paradox? Analysis of a model system with both competition and conmiensalism. Mar Ecol. Prog. Ser 25, 23-30. Bronk, D. A. (2002). Dynamics of DON. In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. Hansell and C. A. Carlson, Eds.), pp. 153-247. Academic Press, San Diego. Bronk, D. A., and Gilbert, P. M. (1993). Application of a 15N tracer method for the measurement of dissolved organic nitrogen release by phytoplankton. Mar Biol. 115, 501-508.

Production and Removal Processes

141

Bronk, D. A., and Ward, B. B. (2000). Magnitude of DON release relative to gross nitrogen uptake in marine systems. Limnol Oceanogr., in press. Brophy, J. E., and Carlson, D. J. (1989). Production of biologically refractory dissolved organic carbon by natural seawater microbial populations. Deep-Sea Res 36,497-507. Bumey, C. M., Johnson, K. M., Lavoie, D. M., and Sieburth, J. M. (1979). Dissolved carbohydrate and microbial ATP in the North Atlantic concentrations and interactions. Deep Sea Res. 26A, 1267-1290. Campbell, L., NoUa, H. A., and Vaulot, D. (1994). The importance of Prochlorophytes to community structure in the central North Pacific Ocean. Limnol. Oceanogr. 39, 954-961. Carlson, C. A., Bates, N. R., Ducklow, H. W., and Hansell, D. A. (1999). Estimation of bacterial respiration and growth efficiency in the Ross Sea, Antarctica. Aquat. Microb. Ecol. 19, 229244. Carlson, C. A., and Ducklow, H. W. (1995). Dissolved organic carbon in the upper ocean of the central Equatorial Pacific, 1992: Daily and fine-scale vertical variations. Deep Sea Res. 7/42,639656. Carlson, C. A., and Ducklow, H. W. (1996). Growth of bacterioplankton and consumption of dissolved organic carbon in the Sargasso Sea. Aquat. Microb. Ecol. 10,69-85. Carlson, C. A., Ducklow, H. W., Hansell, D. A., and Smith, W. O. (1998). Organic carbon partitioning during spring phytoplankton blooms in the Ross Sea Polynya and the Sargasso Sea. Limnol. Oceanogr 43. Carlson, C. A., Ducklow, H. W., and Michaels, A. F. (1994). Annual flux of dissolved organic carbon from the euphotic zone in the Northwestern Sargasso Sea. Nature 371,405^08. Carlson, C. A., Ducklow, H. W., and Sleeter, T. D. (1996). Stocks and dynamics of bacterioplankton in the northwestern Sargasso Sea. Deep Sea Res. 1143,491-515. Carlson, C. A., Giovannoni, S. J., Hansell, D. A., Goldberg, S. J., Parsons, R., Otero, M. P., Vergin, K., and Wheeler, B. R. (in press). The effect of nutrient amendments on bacterioplankton growth, DOC utilization, and community structure in the northwestern Sargasso Sea. Aquat. Microb. Ecol. Carlson, C. A., Hansell, D. A., Peltzer, E. T., and Smith, W O. (2000). Stocks and dynamics of dissolved and particulate organic matter in the southern Ross Sea, Antarctica. Deep Sea Res. II47, 3201-3225. Carlson, C. A., Hansell, D. A., and Smith, W. O. (in press). Filtration volume and retention of carbon-14 labeled particulate organic carbon on GF/F filters during a Phaeocystis antarctica bloom in the Ross Sea, Antarctica. Antarct. J. US. Carlson, D. J., Mayer, L. M., Brann, M. L., and Mague, T. H. (1985). Binding of monomeric organic compounds to macromolecular dissolved organic matter in seawater. Mar Chem 16,141-153. Caron, D. A. (2000). Symbiosis and mixotropy among pelagic microorganisms. In "Microbial Ecology of the Oceans" (D. L. Kirchman, Ed.), pp. 495-523. Wiley-Liss, New York. Caron, D. A., Goldman, J. C , Andersen, O. K., and Dennett, M. R. (1985). Nutrient cycling in a microflagellate food chain. II. Population dynamics and carbon cycling. Mar Ecol. Prog. Ser 24, 243-254. Caron, D. A., Lim, E. L., Miceli, G., Waterbury, J. B., and Valois, F. W. (1991). Grazing and utilization of chroococcoid cyanobacteria and heterotrophic bacteria by protozoa in laboratory cultures and a coastal plankton community. Mar Ecol. Prog. Ser 76,205-217. Caron, D. A., Dam, H. G., Kremer, P., Lessard, E. J., Madin, L. P., Malone, T. C , Napp, J. M., Peele, E. R., Roman, M. R., and Youngbluth, M. J. (1995). The contribution of microorganisms to particulate carbon and nitrogen in surface waters of the Sargasso Sea near Bermuda. Deep Sea Res. 142,943-912. Caron, D. A., Lin, E. L., Sanders, R. W, Dennett, M. R., and Bemiger, U. G. (2000). Response of bacterioplankton and phytoplankton to organic carbon and inorganic nutrient additions in contrasting oceanic ecosystems. Aquat. Microb. Ecol. 22,175-184.

142

Craig A. Carlson

Cherrier, J., Bauer, J. E., and Druffel, E. R. M. (1996). Utilization and turnover of labile dissolved organic matter by bacterial heterotrophs in eastern North Pacific surface waters. Mar. Ecol. Prog. Sen 139,267-279. Cherrier, J., Bauer, J. E., Druffel, E. R. M., Coffin, R. B., and Chanton, J. R (1999). Radiocarbon in marine bacteria: Evidence for the ages of assimilated carbon. Limnol. Oceanogr. 44,730-736. Chin-Leo, G., and Benner, R. (1992). Enhanced bacterioplankton production and respiration at intermediate salinities in the Mississippi River plume. Mar. Ecol. Prog. Sen 87, 87-103. Chisholm, S. W., Olson, R. J., Zettler, E. R., Goericke, R., Waterbury, J. B., and Welschmeyer, N. A. (1988). A novel free-living prochlorophyte abundant in the oceanic euphotic zone. Nature 334, 340-343. Cho, B. C , and Azam, F. (1988). Major role of bacteria in biogeochemical fluxes in the ocean's interior. Nature 332,441^^3. Christian, J. R. (1995) "Biochemical Mechanisms of Bacterial Utilization of Dissolved and Particulate Organic Matter in the Upper Ocean." Ph.D. dissertion, University of Hawaii. Christian, J. R., and Karl, D. M. (1995). Bacterial ectoenzymes in marine waters: Activity ratios and temperature responses in three oceanographic provinces. Limnol. Oceanogr. 40,1042-1049. Cochlan, W. P., Wikner, J., Steward, G. R, Smith, D. C , and Azam, F. (1993). Spatial distribution of viruses, bacteria and chlorophyll a in neritic, oceanic and estuarine environments. Mar Ecol. Prog. Sen 92,77-87. Coffin, R. B., Connolly, J. P., and Harris, P. S. (1993). Availability of dissolved organic carbon to bacterioplankton examined by oxygen utilization. Mar Ecol. Prog. Sen 101,9-22. Cole, J. J., Findlay, S., and Pace, M. L. (1988). Bacterial production in fresh and saltwater ecosystems: A cross-system overview. Mar. Ecol. Prog. Ser 43, 1-10. Copin-Montegut, G., and Avril, B. (1993). Vertical distribution and temporal variation of dissolved organic carbon in the North-Western Mediterranean Sea. Deep Sea Res. 40,1963-1972. Copping, A. E., and Lorenzen, C. J. (1980). Carbon budget of marine phytoplankton-herbivore system with carbon-14 as tracer. Limnol. Oceanogr 25, 873-882. Cotner, J. B., Ammerman, J. W., Peele, E. R., and Bentzen, E. (1997). Phosphorus-limited bacterioplankton growth in the Sargasso Sea. Aquat. Microb. Ecol. 13,141-149. Cowie, G. L., and Hedges, J. I. (1994). Biochemical indicators of diagenetic alteration in natural organic matter mixtures. Nature 369, 304-307. Dagg, M. J. (1976). Complete carbon and nitrogen budget for the carnivorous amphipod, Calliopus laeviusculus (Kroyer). Int. Rev. gesamten Hydrobiol. 61, 297-357. Dagg, M., Cowles, T., Whitledge, T., Smith, S., Howe, S., and Judkin, D. (1980). Grazing and excretion by zooplankton in the Peru upweUing system during April 1977. Deep Sea Res. 27,43-59. Dafner, E. V. (1992). Dissolved organic carbon in waters of the Polar Frontal Zone of the Atlantic Antarctic in the spring-summer season of 1988-1989. Mar Chem. 37, 275-283. Daley, K. L., Wallace, D. W R., Smith, W. O., Skoog, A., Lara, R., Gosselin, M., Falck, E., and Yager, P. L. (1999). Non-Redfield carbon and nitrogen cycling in the Arctic: Effects of ecosystem structure and dynamics. J. Geophys. Res. 104, 3185-3199. Davidson, A. T., and Marchant, H. J. (1992). Protist abundance and carbon concentration during a Phaeocystis-dominated bloom at an Antarctic coastal site. Polar Biol. 12, 387-395. Decho, A. W. (1990). Microbial exopolymer secretions in ocean environments: Their role(s) in food webs and marine processes. Oceanogr Mar Biol. Annu. Rev. 28,73-153. del Giorgio, P., and Cole, J. J. (1998). Bacterial growth efficiency in natural aquatic systems. Annu. Rev. Ecol. Syst. 29,503-541. del Giorgio, P. A., and Cole, J. J. (2000). Bacterial energetics and growth efficiency. In "Microbial Ecology of the Oceans" (D. L. Kirchman, Ed.), pp. 289-325. Wiley-Liss, New York, del Giorgio, P. A., Cole, J. J., andCibleris, A. (1997). Respiration rates in bacteria exceed phytoplankton production in unproductive aquatic systems. Nature 385,148-151

Production and Removal Processes

143

DeLong, E. R, Wu, K. Y, Prezelin, B. B., and Jovine, R. V. M. (1994). High abundance of Archaea in Antarctic marine picopiankton. Nature 371,695-697. Druffel, E. R. M., Bauer, J. E., Williams, R M., Griffin, S., and Wolgast, D. (1996). Seasonal variability of particulate organic radiocarbon in the northeast Pacific Ocean. /. Geophys. Res. 101, 20, 54320, 552. Druffel, E. R. M., Griffin, S., Bauer, J. E., Wolgast, D. M., and Wang, X.-C. (1998). Distribution of particulate organic carbon and radiocarbon in the water column from the upper slope to the abyssal NE Pacific Ocean. Deep Sea Res. II45, 667-687. Druffel, E. R. M., Williams, R M., Bauer, J. E., and Ertel, J. R. (1992). Cycling of dissolved and particulate organic matter in the open ocean. /. Geophys. Res. 97,15,639-15,659. Ducklow, H. W. (1999). The bacterial component of the oceanic euphotic zone. FEMS Microbiol. Ecol. 30,1-10. Ducklow, H. W (2000). Bacterial production and biomass in the ocean. In "Microbial Ecology of the Oceans" (D. L. Kirchman, Ed.), pp. 85-120. Wiley-Liss, New York. Ducklow, H. W, and Carlson, C. A. (1992). Oceanic Bacterial Production. In "Advances in Microbial Ecology" (K. C. Marshall, Ed.), Vol. 12, pp. 113-181. Plenum Press, New York. Ducklow, H. W, Carlson, C. A., Bates, N. R., Knap, A. H., and Michaels, A. E (1995). Dissolved organic carbon as a component of the biological pump in the North Atlantic Ocean. Phil.Trans. R. Soc. A 348,161-167. Ducklow, H. W, Carlson, C. A., Church, M., Kirchman, D. L., Smith, D. C , and Steward, G. (2001). The seasonal development of the bacterioplankton bloom in the Ross Sea, Antarctica, 1994-1997. Deep Sea Res. II, 48,4199^221. Ducklow, H. W, Dickson, M. L., Kirchman, D. L., Steward, G., Ochardo, J., Marra, J., and Azam, F. (2001). Constraining bacterial production conversion efficiency and respiration in the Ross Sea, Antarctica, January-February, 1997. Deep Sea Res. 7/47:3227-3248. Ducklow, H. W, Purdie, D. A., WiUiams, P. J. LeB, and Davies, J. M. (1986). Bacterioplankton: A sink for carbon in a coastal plankton community. Science 232:865. Duursma, E. K. (1963). The production of dissolved organic matter in the sea, as related to the primary gross production of organic matter. Neth. J. Sea Res. 2, 85-94. Fasham, M. J. E, Balino, B., Bowles, M., et al, (2001). Ocean Biogeochemistry and Climate Change: JGOFS Achievements 1988-2000. Ambio (Stockholm), 10,4-31. Ferrier-Pages, C , Kamer, M., and Rassoulzadegan, F. (1998). Release of dissolved amino acids by flagellates and ciliates grazing on bacteria. Oceanol. Acta 21,485-494. Floodgate, G.D. (1995). Some environmental aspects of marine hydrocarbon bacteriology. Aquat. Microb. Ecol. %3-n. Fogg, G. E. (1966). The extracellular products of algae. Oceanogn Mar. Biol. Annu. Rev. 4,195-212. Fogg, G. E. (1983). The Ecological significance of extracellular products of phytoplankton. Bot. Mar 26, 3-14. Fuhrman, J. A. (1992). Bacterioplankton roles in cycling of organic matter: The microbial food web. In "Primary Productivity and Biogeochemical Cycles in the Sea" (P. G. Falkowski and A. D. Woodhead, Ed.), pp. 361-383. Plenum Press, New York. Fuhrman, J. A. (1999). Marine viruses and their biogeochemical and ecological effects. Nature 399, 541-548. Fuhrman, J. A., and Ferguson, R. L. (1986). Nanomolar concentrations and rapid turnover of dissolved free amino acids in seawater: Agreement between chemical and microbiological measurements. Mar Ecol. Prog. Ser 33, 237-242. Giovannoni, S. J., Rappe, M. S., Vergin, K., and Adair, N. (1996). 16S rRNA genes reveal stratified open ocean bacterioplankton populations related to the green non-sulfur bacteria phylum. Proc. Natl. Acad. Sci. USA 93,7979-7984. Gobler, C. J., Hutchins, D. A., Fisher, N. S., Cosper, E. M., and Saiiudo-Wilhelmy, S. A. (1997).

144

Craig A. Carlson

Release and bioavailability of C, N, P, Se and Fe following viral lysis of marine chrysophyte. Limnol. Oceanogr. 42, 1492-1504. Goldman, J. C , and Dennett, M. R. (1985). Susceptibility of some marine phytoplankton species to cell breakage during filtration and post-filtration rinsing. J. Exp. Mar. Biol. Ecol. 86,47-58. Goldman, J. C , Hansell, D. A., and Dennett, M. R. (1992). Chemical characterization of three large oceanic diatoms: Potential impact on water column chemistry. Mar. Ecol. Prog. Sen 88,257-270. Goldman, J. C , and Dennett, M. R. (2000). Growth of marine bacteria in batch and continuous culture under carbon and nitrogen limitation. Limnol. Oceanogr 45,789-800. Gordon, D. A., and Giovannoni, S. J. (1996). Stratified microbial populations related to Chorobium and Fibrobacter detected in the Atlantic and Pacific oceans. Appl. Environ. Microbiol. 62,1171-1177. Gosselin, M., Levasseur, M., Wheeler, P. A., Homer, R. A., and Booth, B. C. (1997). New measurements of phytoplankton and ice algal production in the Arctic Ocean. Deep Sea Res. II44,16231644. Griffith, P. C , Douglas, D. J., and Wainright, S. C. (1990). MetaboHc activity of size-fractionated microbial plankton in estuarine, near-shore, and continental shelf waters of Georgia. Mar Ecol. Prog. Ser 59, 263-270. Guillard, R. R. L., and Hellebust, J. A. (1971). Growth and the production of extracellular substances by two strains of Phaeocystis Pouchetii. J. Phycol. 7, 330-338. Guillard, R. R. L., and Wangersky, P. J. (1958). The production of extracellular carbohydrates by some marine flagellates. Limnol. Oceanogr 3,449^54. Hagstrom, A., Azam, R, Andersson, A., Wikner, J., and Rassoulzadegan, F. (1988). Microbial loop in an oligotrophic pelagic marine ecosystem: Possible roles of cyanobacteria and nanoflagellate in the organic fluxes. Mar Ecol. Prog. 5^r. 49,171-178. Hansell, D. A. (2001). DOC in the global ocean carbon cycle. In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. Hansell and C. A. Carlson, Ed.), pp. 685-715. Academic Press, San Diego. Hansell, D. A., Bates, N. R., and Gundersen, K. (1995). Mineralization of dissolved organic carbon in the Sargasso Sea. Mar Chem. 51, 201-212. Hansell, D. A., and Carlson, C. A. (1998a). Deep ocean gradients in dissolved organic carbon concentrations. Nature 395, 263-266. Hansell, D. A., and Carlson, C. A. (1998b). Net community production of dissolved organic carbon. Global Biogeochem, Cycles 12,443^53. Hansell, D. A., and Carlson, C. A. (2001). Biogeochemistry of total organic carbon and nitrogen in the Sargasso Sea: Control by convective overturn. Deep Sea Res. II4Sj 1649-1667. Hansell, D. A., Carlson, C. A., and Suzuki, Y. (2002). Dissolved organic carbon export with North Pacific intermediate water formation. Global Biogeochem. Cycles, in press. Hansell, D. A., and Waterhouse, T. Y. (1997). Controls on the distribution of organic carbon and nitrogen in the eastern Pacific Ocean. Deep-Sea Res. 144, 843-857. Hanson, L. J., and Norrman, B. (1995). Release of dissolved organic carbon (DOC) by the scyphozoan jellyfish Aurelia aurita and its potential influence on the production of planktonic bacteria. Mar Biol. 121, 527-532. Harvey, G. R., Boran, D. A., Chesal, L. A., and Tokar, J. M. (1983). The structure of marine fulvic and humic acids. Mar Chem. 12, 119-132. Hedges, J. I. (1988). Polymerization of humic substances in natural environments. In "Polymerization of Humic Substances in Natural Environments" (F. H. Frimmel and R. F. Christman, Eds.), pp. 4 5 ^ 8 . Wiley, Chichester. Heissenberger, A., Leppard, G. G., and Hemdl, G. J. (1996). Relationship between the intracellular integrity and the morphology of the capsular envelope in attached and free-living marine bacteria. Appl. Environ. Microbiol. 62,4521-4528. Hellebust, J. A. (1965). Excretion of some organic compounds by marine phytoplankton. Limnol. Oceanogr 10, 192-206.

Production and Removal Processes

145

Hoppe, H. G. (1991). Microbial extracellular enzyme activity: A new key parameter in aquatic ecology. In "Microbial Enzymes in Aquatic Environments" (R. J. Chrost, Ed.), pp. 60-83. Springer-Verlag, New York. Hu, S., and Smith, W. O. (1998). The effects of irradiance on nitrate uptake and dissolved organic nitrogen release by phytoplankton in the Ross Sea. Cont. Shelf Res. 18,971-990. Ikeda, T., and Skjoldal, H. R. (1989). Metabolism and elemental composition of zooplankton from the Barents Sea during early Arctic sununer. Mar. Biol. 100,173-183. Imai, I., Ishida, Y., and Hata, Y. (1993). Killing of marine phytoplankton by a gliding bacterium Cytophaga sp., isolated from the coastal sea of Japan. Mar. Biol. 116, 527-532. Ittekkot, v., Brockmann, U., Michaelis, W., and Degens, E. T. (1981). Dissolved free and combined carbohydrates during a phytoplankton bloom in the northern North Sea. Mar Ecol. Prog. Ser 4, 299-305. Jahnke, R., and Craven, D. B. (1995). Quantifying the role of heterotrophic bacteria in the carbon cycle: A need for respiration rate measurements. Limnol. Oceanogr 40,436-441. Jawed, M. (1969). Body nitrogen and nitrogenous excretion in Neomysis rayii Murdoch and Eupausia pacific Hansen. Limnol. Oceanogr 14,748-754. Jenkins, W. (1980). Tritium and 3He in the Sargasso Sea. / Mar Res. 38, 533-569. Johnson, K. M., Sieburth, J. M., WilHams, R J. leb., and Brandstrom, L. (1987). Coulometric total carbon dioxide analysis for marine studies: Automation and calibration. Mar Chem. 21, 117133. J0rgensen, N., O. G., Kroer, N., and Coffin, R. B. (1994). Utilization of dissolved nitrogen by heterotrophic bacterioplankton: Effect of substrate C/N ratio. Appl. Environ. Microbiol. 60, 41244133. Jumars, P. A., Penry, D. L., Baross, J. A., Perry, M. J., and Frost, B. W. (1989). Closing the microbial loop: Dissolved carbon pathway to heterotrophic bacteria from incomplete ingestion, digestion and absorption in animals. Deep-Sea Res. 36,483^95. Kadurugamuwa, J. L,, and Beveridge, T. J. (1996). Bacteriolytic effect of membrane vesicles from Pseudomonas aeruginosa on other bacteria including pathogens: Conceptually new antibiotics. /. Bacteriol. 178, 2161-211 A. Kahler, P., Bj0msen, P. K., Lochte, K., and Anita, A. (1997). Dissolved organic matter and its utiUzation by bacteria during spring in the Southern Ocean. Deep Sea Res. II44, 341-353. Kahler, P., and Koeve, W. (2001). Dissolved organic matter in the sea: Can its C:N ratio explain carbon overconsumption? Deep Sea Res. 148,49-62. Karl, D. M., Rebel, D. V., Bjorkman, K., and Leteher, R. M. (1998). The role of dissolved organic matter release in the productivity of the oligotrophic North Pacific Ocean. Limnol. Oceanogr 43, 1270-1286. Kamer, M. B., Delong, E. P., and Karl, D. M. (2001). Archaeal dominance in the mesopelagic zone of the Pacific Ocean. Nature 409; 507-510. Kamer, M., and Hemdl, G. J. (1992). Extracellular enzymatic activity and secondary production in free-living and marine-snow-associated bacteria. Mar Biol. 113; 341-347. Keil, R. G., and Kirchman, D. L. (1991). Contribution of dissolved free amino acids and ammonium to the nitrogen requirements of heterotrophic bacterioplankton. Mar Ecol. Prog. Ser 73,1-10. Keil, R. G., and Kirchman, D. L. (1994). Abiotic transformation of labile protein to refractory protein in sea water. Mar Chem. 45,187-196. Keil, R. G., and Kirchman, D. L. (1999). Utilization of dissolved protein and amino acids in the northern Sargasso Sea. Aquat. Microb. Ecol. 18,293-300. Kepkay, P. E., Jellett, J. P., and Niven, S. E. H. (1997). Respiration and the carbon-to-nitrogen ratio of a phytoplankton bloom. Mar Ecol. Prog. Ser 150, 249-261. Kieber, D. J., McDaniel, J., and Mopper, K. (1989). Photochemical source of biological substrates in sea water: ImpHcations for carbon cycling. Nature 341,637-639.

146

Craig A. Carlson

Kirchman, D. L. (1990). Limitation of bacterial growth by dissolved organic matter in the subarctic Pacific. Mar. Ecol Prog. Ser 62,47-54. Kirchman, D. L. (2000). Uptake and regeneration of inorganic nutrients by marine heterotrophic bacteria. In "Microbial Ecology of the Oceans" (D. L. Kirchman, Ed.), pp. 261-288. Wiley-Liss, New York. Kirchman, D. L., Henry, D. L., and Dexter, S. C. (1989). Adsorption of protein to surfaces in seawater. Mar Chem. 27,201-211. Kirchman, D.L., Keil, R., and Wheeler, P. A. (1990). Carbon limitation of ammonium uptake by heterotrophic bacteria in the subarctic Pacific. Limnol. Oceanogn 35,1267-1278. Kirchman, D.L., Keil, R. G., and Wheeler, P. A. (1989). The effect of amino acids on ammonium utilization and regeneration by heterotrophic bacteria in the subarctic Pacific. Deep-Sea Res 36, 1763-1776. Kirchman, D. L., Lancelot, C , Fasham, M., Legendre, L., Radach, G., and Scott, M. (1993). Dissolved organic matter in biogeochemical models of the ocean. In "Towards a Model of Ocean Biogeochemical Processes" (G. T. Evans and M. J. R. Fasham, Ed.), pp. 209-225. Springer-Verlag, Berlin Heidelberg. Kirchman, D. L., Meon, B., Ducklow, H. W, Carlson, C. A., Hansell, D. A., and Steward, G. (2001). Glucose fluxes and concentrations of dissolved combined neutral sugars (polysaccharides) in the Ross Sea and Polar Front Zone, Antarctica. Deep Sea Res. II, 48,4179-4197. Kirchman, D. L., Suzuki, Y., Garside, C , and Ducklow, H. W. (1991). High turnover rates of dissolved organic carbon during a spring phytoplankton bloom. Nature 352,612-614. Kremer, P. (1977). Respiration and excretion by Ctenophore Mnemiopsis leidyi. Mar. Biol. 44,43-50. Kroer, N. (1993). Bacterial growth efficiency on natural dissolved organic matter. Limnol. Oceanogr. 38,1282-1290. Lampert, W (1978). Release of dissolved organic carbon by grazing zooplankton. Limnol. Oceanogn 23,831-834. Lampitt, R. S., Noji, T, and Bodugen, B. v. (1990). What happens to zooplankton fecal pellets? Imphcation for material flux. Mar Biol. 104, 15-23. Lancelot, C. (1979). Gross excretion rates of natural marine phytoplankton and heterotrophic uptake of excreted products in the southern North Sea, as determined by short-term kinetics. Mar Ecol. Prog. Ser 1, 179-186. Lancelot, C. (1983). Factors affecting phytoplankton extracellular release in the Southern Bight of the North Sea. Mar Ecol. Prog. Ser 12, 115-121. Lancelot, C , and Billen, G. (1984). Activity of heterotrophic bacteria and its coupling to primary production during spring phytoplankton bloom in the southern bight of the North Sea. Limnol. Oceanogr 29,121-130. Lara, R. J., and Thomas, D. N. (1995). Formation of recalcitrant organic matter: Humification dynamics of algal derived dissolved organic carbon and its hydrophobic fractions. Mar Chem. 51,193-199. Larsson, U., and Hagstrom, A. (1979). Phytoplankton exudate release as an energy source for the growth of pelagic bacteria. Mar Biol 52, 199-206. Larsson, U., and Hagstrom, A. (1982). Fractionated phytoplankton primary production, exudate release and bacterial production in a Baltic eutrophication gradient. Mar Biol 67, 57-70. Lasker, R. (1966). Feeding growth and respiration and carbon utilization of euphausiid crustacean. /. Fish. Res. Bd. Can. 23, 1291-1317. Lee, D. Y, and Rhee, G. Y (1997). Kinetics of cell death in the cyanohactenum Anabaena flos-auae and the production of dissolved organic carbon. J. Phycol. 33,991-998. Legendre, L., and Le Fevre, J. (1995). Microbial food webs and the export of biogenic carbon in oceans. Aquat. Microb. Ecol. 9, 69-77. Lessard, E. J., Foy, M. S., Garrison, D. L., and Gowing, M. M. (1995). Grazing on phytoplankton blooms in the Ross Sea polynya: November and December 1994. Antarct. J. US 30, 214-215.

Production and Removal Processes

147

Lewin, R. A. (1956). Extracellular polysaccharides of green algae. Can. J. Microbiol. 2,665-672. Lignell, R. (1990). Excretion of organic carbon by phytoplankton: Its relation to algal biomass, primary productivity and bacterial secondary productivity in the Baltic Sea. Mar. Ecol. Prog. Sen 68,85-99. Liu, H., NoUa, H. A., and Campbell, L. (1997). Prochlorococcus growth rate and contribution to primary production in the equatorial and subtropical North Pacific Ocean. Aquat. Microb. Ecol. 25, 39-47. Longhurst, A. R., and Harrison, W. G. (1989). The biological pump: Profiles of plankton production and consumption in the upper ocean. Prog. Oceanogr 22,47-123. MacKenzie, F. T. (1981). Global carbon cycle: Some minor sinks for CO2. In "Flux of Organic Carbon by Rivers to the Ocean" (G. E. Likens, F. T. MacKenzie, J. E. Richey, J. R. Sedell, and K. K. Turekian, Ed.), pp. 360-384. U.S. Dept. of Energy, Washington, DC. Mague, T. H., Freberg, E., Hughes, D. J., and Morris, L (1980). Extracellular release of carbon by marine phytoplankton; A physiological approach. Limnol. Oceanogr 25,262-279. Manahan, D. T., and Richardson, K. (1983). Competition studies on the uptake of dissolved organic nutrients by bivalve larvae (Mytilus edulis) and marine bacteria. Mar Biol. 75, 241-247. Marchant, H. J., and Scott, F. J. (1993). Uptake of sub-micrometer particles and dissolved organic material by Antarctic choanoflagellates. Mar Ecol. Prog. Ser 92,59-64. Mayzaud, P. (1973). Respiration and nitrogen excretion of zooplankton. II. Studies of the metabolic characteristics of starved animals. Mar Biol. 21, 19-28. McCarthy, M., Hedges, J., and Benner, R. (1996). Major biochemical composition of dissolved high molecular weight organic matter in seawater. Mar Chem. 55, 281-297. McCarthy, M. D., Hedges, J. I., and Benner, R. (1998). Major bacterial contribution to marine dissolved organic nitrogen. Science 281, 231-234. Michaels, A. F , and Silver, M. W. (1988). Primary production, sinking fluxes and the microbial food web. Deep Sea Res. 35,473^90. Middelboe, M., J0rgensen, N. O. G., and Kroer, N. (1996). Effects of viruses on nutrient turnover and growth efficiency of noninfected marine bacterioplankton. Appl. Environ. Microbiol. 62, 19911997. M0ller-Jensen, L. (1983). Phytoplankton release of extracellular organic carbon, molecular weight composition, and bacterial assimilation. Mar Ecol. Prog. Ser 11, 3 9 ^ 8 . Mopper, K., and Kieber, D. J. (2002). Photochemistry and cycling of carbon, sulfur, nitrogen and phosphorus. In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. Hansell and C. A. Carlson, Eds.), pp. 455-507. Academic Press, San Diego. Mopper, K., Zhou, X. L., Kieber, R. J., Kieber, D. J., Sikorski, R. J., and Jones, R. D. (1991). Photochemical degradation of dissolved organic carbon and its impact on the oceanic carbon cycle. Nature 353,60-62. Moran, M. A., and Zepp, R. G. (1997). Role of photoreactions in the formation of biologically labile compounds from dissolved organic matter. Limnol. Oceanogr 42, 1307-1316. Moran, M. A., and Zepp, R. G. (2000). UV radiation effects on microbes and microbial processes. In "Microbial Ecology of the Oceans," pp. 201-228. Wiley-Liss, New York. Myklestad, S. O., Holm-Hansen, O., Varum, K. M., and Volcani, B. E. (1989). Rate of release of extracellular amino acids and carbohydrates from marine diatom Chatoceros affinis. J. Plankton Res. 11,763-773. Naganuma, T., Konishi, S., Inoue, T., Nakane, T., and Sukizaki, S. (1996). Photodegradation of photoalteration? Microbial assay of the effect of UV-B on dissolved organic matter. Mar Ecol. Prog. 5^^135,309-310. Nagata, T. (2000). Production mechanisms of dissolved organic matter. In "Microbial Ecology of the Oceans" (D. L. Kirchman, Ed.), pp. 121-152. Wiley-Liss, New York. Nagata, T., and Kirchman, D. L. (1991). Release of dissolved free and combined amino acids by bacterivorous marine flagellates. Limnol. Oceanogr 36,433^43.

148

Craig A. Carlson

Nagata, T., and Kirchman, D. L. (1992). Release of dissolved organic matter by heterotrophic protozoa: Implications for microbial foodwebs. Arch. Hydrobiol. 35,99-109. Nagata, T., and Kirchman, D. L. (1996). Bacterial degradation of protein adsorbed to model submicron particles in seawater. Mar. Ecol Prog. Ser. 132,241-248. Norrman, B., Zweifel, U. L., Hopkinson, C. S., and Fry, B. (1995). Production and utilization of dissolved organic carbon during an experimental diatom bloom. Limnol. Oceanogr. 40, 898-907. Obemoster, I., and Hemdl, G. J. (1995). Phytoplankton extracellular release and bacterial growth: Dependence on the inorganic N:P ratio. Man Ecol. Prog. Sen 116,247-257. Ogawa, H., Fukuda, R., and Koike, I. (1999). Vertical distribution of dissolved organic carbon and nitrogen in the Southern Ocean. Deep Sea Res. 746,1809-1826. Ogawa, H., Amagai, Y., Koike,!., Kaiser, K., and Benner, R. (2001). Production of refractory dissolved organic matter by bacteria. Science 292,917-920. Ogura, N. (1972). Rate and extent of decomposition of dissolved organic matter in the surface water. Mar Biol. 13, 89-93. Olsen, G. J. (1994). Archaea, Archaea everywhere. Nature 371,657-658. Otake, T., Nogami K., and Maruyama, K. (1993). Dissolved and particulate organic matter as possible food sources for eel leptocephali. Mar Ecol. Prog. Ser 92,27-34. Ouvemey, C. C , and Fuhrman, J. A., (20(X)). Marine planktonic Archaea take up amino acids. Appl. Environ. Microbiol 66,4829-4833. Pakulski, J. D., and Benner, R. (1994). Abundance and distribution of carbohydrates in the ocean. Limnol. Oceanogr 39, 930-940. Parsons, T. R., Le Brasseur, R. J., and Barraclough, W. E. (1970). Levels of production in the pelagic environment of the Strait of Georgia, British Columbia: A review. J. Fish, Res. Bd. Can. 27,12511264. Pomeroy, L. R. (1974). The ocean's food web, a changing paradigm. BioScience 24,499-504. Pomeroy, L. R., and Deibel, D. (1986). Temperature regulation of bacterial activity during the spring bloom in Newfoundland coastal waters. Science 233, 359-361. Pomeroy, L. R., Sheldon, J. E., Sheldon, W. M., and Peters, F. (1995). Limits to growth and respiration of bacterioplankton in the Gulf of Mexico. Mar Ecol. Prog. Ser 111, 259-268. Pomeroy, L. R., Wiebe, W. J., Deibel, D., Thompson, R. J., Rowe, G. R., and Pakulski, J. D. (1991). Bacterial response to temperamre and substrate concentrations during the Newfoundland spring bloom. Mar Ecol. Prog. Ser 75, 143-159. Proctor, L. M., and Fuhrman, J. A. (1990). Viral mortality of marine bacteria and cyanobacteria. Nature 343, 60-62. Quay, R D., Tilbrook, B., and Wong, C. S. (1992). Oceanic uptake of fossil fuel CO2: Carbon-13 evidence. Science 256,74-79. Rich, J., Gossehn, M., Sherr, E., Sherr, B., and Kirchman, D. L. (1997). High bacterial production, uptake and concentrations of dissolved organic matter in the Central Arctic Ocean. Deep Sea Res. 7/44,1645-1663. Rich, J. H., Ducklow, H. W., and Kirchman, D. L. (1996). Concentration and uptake of neutral monosaccharides along 140°W in the equatorial Pacific: Contribution of glucose to heterotrophic bacterial activity and the DOM flux. Limnol. Oceanogr 41,595-604. Rivkin, R. B., and Anderson, M. R. (1997). Inorganic nutrient limitation of oceanic bacterioplankton. Limnol. Oceanogr 42, 730-740. Rivkin, R. B., and Putt, M. (1987). Heterotrophy and photoheterotrophy by Antarctic microalgae: Light-dependent incorporation of amino acids and glucose. J. Phycol. 23, AA2-A52. Sharp, J. H. (1977). Excretion of organic matter by marine phytoplankton: Do healthy cells do it? Limnol. Oceanogr 11, 381-399. Sherr, E., and Sherr, B. (1988). Roles of microbes in pelagic food webs: A revised concept. Limnol. Oceanogr 33,1225-1257.

Production and Removal Processes

149

Shiah, F. K., Ducklow, H. W. (1994). Temperature regulation of heterotrophic bacterioplankton abundance, production, and specific growth rate in Chesapeake Bay. Limnol Oceanogr. 39,1243-1258. Skoog, A., Benner, R. (1997). Aldoses in various size fractions of marine organic matter: Implications for carbon cycling. Limnol. Oceanogr. 42,1803-1813. Skoog, A., Biddanda, B., and Benner, R. (1999). Bacterial utilization of dissolved glucose in the upper water column or the Gulf of Mexico. Limnol. Oceanogr. 44,1625-1633. Smith, D., Simon, M., AUdredge, A. L., and Azam, F. (1992). Intense hydrolytic enzyme activity on marine aggregates and implications for rapid particle dissolution. Nature 359,139-142. Smith, D. C , Steward, G. F , Long, R. A., and Azam, F. (1995). Bacterial mediation of carbon fluxed during a diatom bloom in a mesocosm. Deep Sea Res. 1142,75-98. Smith, R. E. (1982). The estimation of phytoplankton production and excretion by carbon-14. Mar Biol. Lett. 3, 325-334. Smith, W. O., Barber, R. T, Hiscock, M. R., and Marra, J. (2000). The seasonal cycle of phytoplankton biomass and primary productivity in the Ross Sea, Antarctica. Deep Sea Res. II47, 3119-3140. Smith, W. O., Barber, R. T, and Huntsman, S. A. (1977). Primary production off the coast of northwest Africa: Excretion of dissolved organic matter and its heterotrophic uptake. Deep Sea Res. 24,35-47. Smith, W. O., Carlson, C. A., Ducklow, H. W., and Hansell, D. A. (1998). Growth dynamics of Phaeocystis antarctic-dominated plankton assemblages from the Ross Sea. Mar Ecol. Prog. Sen 168, 229-244. Smith, W. O., and Gordon, L. I. (1997). Hyperproductivity of the Ross Sea (Antarctica) polynya during austral spring. Geophys. Res. Lett. 24,233-236. S0ndergaard, M., and Middelboe, M. (1995). A cross-system analysis of labile dissolved organic carbon. Mar Ecol. Prog. Ser 118, 283-294. S0ndergaard, M., and Schierup, H. H. (1981). Release of extracellular organic carbon during a diatom bloom in Lake Moss0: Molecular weight fractionation. Freshwater Biol. 12, 313-320. S0ndergaard, M., Williams, R J. le B., Cauwet, G., Riemann, B., Robinson, C , Terzic, S., Woodward, E. M. S., and Worm, J. (2000). Net accumulation and flux of dissolved organic carbon and dissolved organic nitrogen in marine plankton communities. Limnol. Oceanogr 45, 1097-111. Stein, J. L., and Simon, M. I. (1996). Archaeal ubiquity. Proc. Natl. Acad. Sci. USA 93,6228-6230. Steinberg, D. K., Carlson, C. A., Bates, N. R., Goldthwait, S. A., Madin, L. R, and Michaels, A. F. (2000). Zooplankton vertical migration and the active transport of dissolved organic and inorganic carbon in the Sargasso Sea. Deep Sea Res. 747,137-158. Steinberg, D. K., Goldthwait, S. A., and Hansell, D. A. (2001). Zooplankton vertical migration and the active transport of dissolved organic and inorganic nitrogen in the Sargasso Sea. Deep Sea Res., Submitted for publication. Steward, G. F , Smith, D. C , and Azam, F. (1996). Abundances and production of bacteria and viruses in the Chukchi Sea. Mar Ecol. Prog. Ser 131,287-300. Stewart, J. R., and Brown, R. M. (1969). Cytophaga that kills or lyses algae. Science 164,1523-1524. Stoderegger, K., and Hemdl, G. J. (1998). Production and release of bacterial capsular material and its subsequent utilization by marine bacterioplankton. Limnol. Oceanogr 43, 877-884. Strom, S, L., Benner, R., Ziegler, S., and Dagg, M. J. (1997). Planktonic grazers are a potentially important source of marine dissolved organic carbon. Limnol. Oceanogr 42,1364-1374. Suttle, C. A. (1994). The significance of viruses to mortality in aquatic microbial communities. Microb. Ecol. 28, 237-243. Suttle, C. A., Chan, A. M., and Cottrell, M. T. (1990). Infection of phytoplankton by viruses and reduction of primary productivity. Nature 347,467-469. Takahashi, T, Wanninkhof, R. T, Feely, R. A., Weiss, R. F , Chipman, D. W, Bates, N. R., Olafsson, J., Sabine, C. L., and Sutherland, C. S. (1999). Net sea-air CO2fluxover the global ocean: An improved estimate based on air-sea /7CO2 difference. In "Proceeding of the 2nd Symposium on CO2 in the Oceans" (Y. Nojiri, Ed.), pp. 9-15. Tsukuba, Japan.

150

Craig A. Carlson

Tanoue, E., Ishii, M., and Midorikawa, T. (1996). Discrete dissolved and particulate proteins in oceanic waters. LimnoL Oceanogr. 41,1334-1343. Tanoue, E., Nishiyama, S., Kamo, M., and Tsugita, A. (1995). Bacterial membranes: Possible source of a major dissolved protein in seawater. Geochim. Cosmochim. Acta 59,2643-2648. Taylor, G. T, Iturriaga, R., and Sullivan, C. W. (1985). Interactions of bacterivorous grazers and heterotrophic bacteria with dissolved organic matter. Mar. Ecol. Prog. Sen 23,129-141. Thingstad, T. F., Hagstrom, A., and Rassoulzadegan, F. (1997). Accumulation of degradable DOC in surface waters: Is it caused by a malfunctioning microbial loop? Limnol. Oceanogr 42, 398^04. Thingstad, T. F , and Lignell, R. (1997). Theoretical models for the control of bacterial growth rate, abundance, diversity and carbon demand. Aquat. Microb. Ecol. 13, 19 - 27. Thingstad, T. F, and Rassoulzadegan, F. (1995). Nutrient limitation, microbial food webs, and biological C-pumps: Suggested interactions in a P-Umited mediterranean. Mar Ecol. Prog. Ser 117,299-306. Thingstad, T. F , Zweifel, U. L., and Rassoulzadegan, F. (1998). P limitation of heterotrophic bacteria and phytoplankton in the northwest Mediterranean. Limnol. Oceanogr 43, 88-94. Thomas, J. P. (1971). Release of dissolved organic matter from natural populations of marine phytoplankton. Mar Biol. 11, 311-323. Tranvik, L., and Kokalj, S. (1998). Decreased biodegradability of algal DOC due to interactive effects of UV radiation and humic matter. Aquat. Microb. Ecol. 14, 301-307. Tranvik, L. J. (1993). Microbial transformation of labile dissolved organic matter into humic-like matter in seawater. FEMS Microbiol. Ecol. 12, 177-183. Tranvik, L. J., Sherr, E. B., and Sherr, B. F. (1993). Uptake and utiUzation of "colloidal DOM" by heterotrophic flagellates in seawater. Mar Ecol. Prog. Ser 92, 301-309. Turley, C. M. (1993). The effect of pressure on leucine and thymidine incorporation by free-living bacteria and by bacteria attached to sinking oceanic particles. Deep Sea Res. 140,2193-2206. Urban-Rich, J. (1999). Release of dissolved organic carbon from copepod fecal pellets in the Greeland Sea. J. Exp. Mar Biol. Ecol. 232, 107-124. Veldhuis, M. J. W., and Admiraal, W. (1985). Transfer of photosynthetic products in gelatinous colonies of Phaeocystis pouchetii (Haptophyceae) and its effect on the measurement of excretion rate. Mar Ecol. Prog. Ser 26, 301-304. Verity, P. G. (1981). The effects of temperature, irradiance and day length on the marine diatom Leptocylindrus danicus Cleve. II. Excretion. J. Exp. Mar Biol. Ecol. 55, 159-169. Vemet, M., Matrai, P. A., and Andreassen, I. (1998). Synthesis of particulate and extracellular carbon by phytoplankton at the marginal ice zone in the Barents Sea. J. Geophys. Res. 103,1023-1037. Wafar, M. (1984). Seasonal changes of dissolved organic matter (C, N, P) in permanently well mixed temperate waters. Limnol. Oceanogr 29,1127-1132. Wangersky, P. J. (1978). Production of dissolved organic matter. In "Marine Ecology" (O. Kinne, Ed.), Vol. IV, pp. 115- 200. Wiley, New York. Webb, K. L., and Johannes, R. E. (1967). Studies of the release of dissolved free amino acids by marine zooplankton. Limnol. Oceanogr 12, 376-382. Wheeler, P A., Gosselin, M., Sherr, E., Thibault, D., Kirchman, D. L., Benner, R., and Whitledge, T. E. (1996). Active cycling of organic carbon in the central Arctic Ocean. Nature 380,697-699. Wiebe, W. J., and Smith, D. F. (1977). Direct measurement of dissolved organic carbon release by phytoplankton and incorporation by microheterotrophs. Mar Biol. 42, 213-223. Wiebinga, C. J., and de Baar, H. J. W (1998). Determination of the distribution of dissolved organic carbon in the Indian sector of the Southern Ocean. Mar Chem. 61,185-201. Williams, P. J. leB. (1970). Heterotrophic utiUzation of dissolved organic compound in the sea. I. Size distribution of population and relationship between respiration and incorporation of growth substrates. J. Mar Biol. Assoc. UK 50, 859-870. WiUiams, P. J. leB (1981). Incorporation of microheterotrophic processes into the classical paradigm of the planktonic food web. Kiel. Meeresforsch. 5, 1-28.

Production and Removal Processes

151

Williams, P. J. leB (1995). Evidence for the seasonal accumulation of carbon-rich dissolved organic material, its scale in comparison with changes in particulate material and the consequential effect on net C/N assimilation ratios. Man Chem. 51,17-29. WiUiams, P. J. leB. (1990). The importance of losses during microbial growth: Conmientary on the physiology, measurement and ecology of the release of dissolved organic material. Mar. Microb. Food Webs 4,175-206. Williams, P. J. leB . (2000). Heterotrophic bacteria and the dynamics of dissolved organic material. In "Microbial Ecology of the Oceans" (D. L. Kirchman, Ed.), pp. 153- 200. Wiley-Liss, New York. WiUiams, P. J. leB., and Yentsch, C. S. (1976). An examination of photosynthetic production, excretion of photosynthetic products, and heterotrophic utilization of dissolved organic compounds with reference to results from a coastal subtropical sea. Mar. Biol. 35, 3 1 ^ 0 . Williams, P. M., and Druffel, E. R. M. (1987). Radiocarbon in dissolved organic matter in the central North Pacific Ocean. Nature 330,246-248. Wolter, K. (1982). Bacterial incorporation of organic substances released by natural phytoplankton populations. Mar Ecol. Prog. Serl, 287-295. Wommack, K. E., and Colwell, R. R. (2000). Virioplankton: Viruses in aquatic ecosystems. Microb. Molec. Biol. Rev. 64,69-114. Wood, A. M., and Valen, L. M. V. (1990). Paradox lost? On the release of energy rich compounds by phytoplankton. Mar Microb. Food Webs 4,103-116. Zlotnik, I., and Dubinsky, Z. (1989). The effect of light and temperature on DOC excretion by phytoplankton. Limnol. Oceanogr 34, 831-839. Zusheng, L., Clarke, A. J., and Beveridge, R. J. (1998). Gram-negative bacteria produce membrane vesicles which are capable of killing other bacteria. /. Bacteriol. 180, 5478-5483. Zweifel, U. L. (1999). Factors controlling accumulation of labile dissolved organic carbon in the Gulf of Riga. Estuarine. Coastal Shelf Sci. 48, 357-370. Zweifel, U. L., and Hagstrom, A. (1995). Total counts of marine bacteria include a large fraction of non-nucleoid-containing bacteria (ghosts). Appl. Environ. Microbiol. 61,2180-2185. Zweifel, U. L., Norrman, B., and Hagstrom, A. (1993). Consumption of dissolved organic carbon by marine bacteria and demand for inorganic nutrients. Mar Ecol. Prog. Ser 101,23-32.

This Page Intentionally Left Blank

Chapter 5

Dynamics of DON Deborah A. Bronk Virginia Institute of Marine Science, College of William and Mary, Gloucester Point, Virginia I. Introduction II. Concentration and Composition of the DON Pool A. Methods for Measuring DON Concentrations B. DON Distributions and Correlative Relationships between DON and Other Parameters C. Chemical Composition of the DON Pool D. Concentration and Composition of the DON Pool: Research Priorities III. Sources of DON A. Biotic Sources of DON in the Water Column B. Methods for Estimating Biotic DON Release Rates C. Literature Values of DON Release

Rates in Aquatic Environments D. Sources of DON: Research Priorities IV. Sinks for DON A. Heterotrophic versus Autotrophic DON Utilization B. Methods for Estimating Biotic DON Uptake C. Literature Values of DON Uptake in Aquatic Environments D. Photochemical Decomposition as a Sink for DON E. Sinks for DON: Research Priorities V. DON Turnover Times VI. Summary References

I. INTRODUCTION Dissolved organic nitrogen (DON) is that subset of the dissolved organic matter (DOM) pool that contains N. From the perspective of a microorganism, this is where the action is—one-stop shopping for N, carbon (C), and energy. Research into DON, however, has lagged far behind that of the larger dissolved organic carbon (DOC) pool as clearly seen by the C:N ratio of chapters in this volume. This situation is primarily the result of the substantial analytical challenges Biogeochemistry of Marine Dissolved Organic Matter Cop)mght 2002, Elsevier Science (USA). All rights reserved.

153

154

Deborah A. Bronk

inherent in DON research. DON exists in substantially lower concentrations than DOC, multiple chemical analyses are required for a single DON measurement, inorganic N removal is a nightmarish undertaking, and unless you have easy access to a nuclear reactor manufacturing short-lived ^^N, one must be content with labor-intensive stable isotopes rather than the quicker and more sensitive radiotracers. The objectives of this chapter are to review available data specific to DON on the concentration and composition of the pool, to describe recent findings on the sources of DON to aquatic systems, and to survey data on rates and mechanisms of DON uptake and other sinks. An exhaustive review of DON was published by Antia et al (1991). Therefore, this review will focus on work pubUshed largely after 1990 and topics not included in the earlier review. As a subset of the DOM pool, much of the information presented on DOC throughout this volume holds equally true for DON.

11. CONCENTRATION AND COMPOSITION OF THE DON POOL Measurements of DON concentrations have become a routine component of many studies. This section reviews methods for measuring DON and then presents a survey of recent literature values of DON concentrations, relationships between DON and other parameters, data on the chemical composition of the pool, and suggested research priorities for the future. Due to space limitations, DON concentrations in lakes, streams, or groundwater, with some exceptions, are not included. A. METHODS FOR MEASURING DON CONCENTRATIONS Studies of any aspect of DON cycling require first and foremost a reliable method of quantifying DON concentrations with high precision (Bronk et ai, 2000; see Sharp, Chapter 2). To calculate DON concentrations, one must first obtain an accurate total dissolved N (TDN) concentration. The TDN pool consists of an inorganic fraction, composed of ammonium (NH4^), nitrate (NOs"), and nitrite (N02~), and an organic fraction (i.e., DON), the composition of which is largely unknown (see Section II.C). There are presently three methods conmionly used to measure TDN concentrations in aquatic systems: persulfate oxidation (Menzel and Vaccaro, 1964; Sharp, 1973; Valderrama, 1981), ultraviolet oxidation (Armstrong etal, 1966; Armstrong and Tibbitts, 1968), and high-temperature oxidation (Sharp, 1973; Suzuki and Sugimura, 1985). After a TDN concentration has been measured, the sum of the NH4'^ and combined NO3" / NO2" concentrations are subtracted from it, with the residual being defined as DON. This approach is problematic

Dynamics of DON

155

because estimates of DON concentrations have the combined analytical error and uncertainty of three analyses: TDN, NH4+, and combined NOs"/ N02~. The first broad community comparison of the three methods used to measure DON was recently completed (Sharp et al, in press; see Sharp, Chapter 2). It consisted of 29 sets of analyses done on five natural samples. The coefficient of variations for the five samples range from 19 to 46%, with the poorest replication observed on deep ocean samples. No one method emerged as clearly superior. B. DON DISTRIBUTIONS AND CORRELATIVE RELATIONSHIPS BETWEEN D O N AND O T H E R PARAMETERS Here DON concentrations are presented and discussed with respect to global distributions, vertical profiles, seasonal variability, and the link between DON and inorganic N distributions. 1. Concentrations of DON in Aquatic Environments In general, the lowest mean concentrations of DON are found in the deep ocean and the highest mean concentrations are found in rivers (Fig. 1 A). Concentrations in Table I for the surface ocean range from 0.8 to 13 JJLM with a mean of 5.8 ± 2.0 /xM. Note that many open ocean studies present data on total organic N (TON), rather than DON. Most researchers working in oligotrophic waters forego the filtration step because the particulate N (PN) pool is generally so small (

q^J

2^

^^

§ a.

OH

9

ON

^

-J" ^ 1> ON ^ ^

ON ON

ON 1 3

OQ

&0

CO

o

o

^

-H

o

2 I

ON VO

d c> -H -H CO vo

^ ^

c

(U

ed a (U (U CO C/3 T3 T3 C C c« 3 C

Tj-

^ 4^ CO 00

in in

coC/5

ol

13

o :z;

>

1

r-

s. a.

;-i

I

o in

;H

C/3 - S

O ^2;

1>

(D

ON

i>

OH

(D Cin

VH

^

OH

>H

u

^

0\

^

i^

T—I

d -H

in

ON r^'

en -H

CO

o

ON

^

OH

iH

(D

i-i

^ D&.

csi

rt ON ON

1

z PU

ON

t


O ^

00

-H

-H

o ri

-H i/i

QO

"^ OS

-H

2


3

o

^ o

CI.

CN

o

^

J

§

»o

so

00 ON

•§

eu

to :z;

1 cC g •g 1

CO c 5 C3

ffi ffi

g> »n p r^ 00

-H -H

Tt od

ZH

l^

•+^ -H

,-H

^

r^ ^^ r-; r-^ wS u-i

ci, ^ ^ o

OX) OS OS

OS OS


^ en p lo

"^

^3 ^

Cvi

-H

^

ffi

-H

ID

^ >

PLH

ri O

KJ

OH

CO

1 '=1 w

u

O

-H

q -H

«o o

-H

8

r^

ON

8

^

vo

ON

f^

ON

^^ £ ^^

(D pg

S ^ a S C/3 21 ^ •s =J p^ o o C/5

g

OM

O

> o ^ o

ON

O

ON

OO

s

o

c3 >^ o

> q

>

2

ON g ,^ c r) "5 .^ 3

o o

p

"^

ON

'-H

Z

o

cn '-^

^t -H

v£)

ra 2

'-H

ON ^t

d -H

2

ON ON ON

O Pu

ON ON ON

PLH

A

ON

o

o

o A O

ON

ON ON

-H

ON ON

> O

>

O

;:^

5 5

'^,

>ri 2

4

S

>f

2

ON

CO

Q

w"

tS"

w"

^ o .2 »^ .5 ^ .2 »^

f^ c ^< - < £< ^ o--3Z--j»n--HCO

(U

Z

'"t CO

XH'

^

1

(U

ON

e 4 o > u

w ;-( o e e o o

s ss

soc

ffl PQ ffl

< < S

t« 3 D

^ ^§

-H -H-H

VO

^

ON

-H -H -H

NO r t 0 0

«o rn

-fl -H -H q »n o

q ^ o

ffi ffi cu

X X

§ ON ^3 ^d^

>

-H

0)

ON ON

X>

S

OJ

on

3 «

•^

13

ON

oo

>/S

"ON

9r

§; _ 1

(/3

3 3


» a PQ PQ PQ

:^

1 < >—>

ON

ON

in ON 00 00

ON

ON

0

ON

ON

3 bO 3

'"' (« en r
o> o

Oi

pji

Z Z

2 u

o o

(N

Im •§

1 I

:z;

ON

p O

ON

q

O

ON

cd

1 ^s u e ,^ 3 O 00

n d ^ ^ KS Kj c3 «J C PQ PQ PQ PQ

^< < <



»

•'C

>»

U

U

S

« « I ? I ? .5 i§ g,i s.i g-i 8-1 U

41 Tt f^ fS

iH

in

-H 00

-H

«H

•so PQ

§

73

o

^

z

•

ON ON

-H ^.

5 "^ en 00

-H o ^ ^*

ON ON ON

s

PQ

^

-H ^. '"^

(N

O

r-

d -H

o d

:z;

ON

cn

UP

^ Q "Z, Z

+j^

^

CN

2

NO

Q Z

^^

c3

9 > E>> E>>

m u u

^

t-:

d

22

ON ^

>>

z

X

+^

+^

X

+^

+Tf^

K Z

+^

+-f

)z;

O

z

2:

^

:i:

z

>>

X

>^ x^

^

£

D

2i

S

^ < < O Q

^ ^ Q

<


a

1/3

x^ 1/3

2^

«

2^

^

<
N

1OS T-H ( N

^

O)

CD

^

A-

Id ^

5-1

O 7 3 OJ J-i _r-

Jf|

fC U

'Q;

nj

'^ B + -^ -^

< c«"

> +O

0)

+

0

ft O ^-^ «

• c l «:s

-d

a

O

CQ

1 •S D X) ^ ;^ ,c

iS

z b

a 13

00

^

i^-s

Q

S^ON

(u

^ t:^ t: ^ ^Q t:

^ ^ ^ ON ^ Q "3 Q J.

^ c(U 22 < ^

•fe

OO

d ON

d

n

I

e> a>> c«>» c3 ^ ^ ^ •c •c a a. < < Cu

•C

t^

u


1 C3 00

«j

&

o

ON

ON ON

'Q

"^

3

ON ON

ON

"^

'Q

3

5:1

;> .^ s^

^0

" Ji 1 J

L-l•-

in

en"

«^ ^ ^ ^ 'Z

en

14

^ :z;

o

ren

-H

8 o

0

.52 3

^3 u

1°

I 1^

>

s>

04

Tt

^ en

/-< in

^"^ Z

-H m en

>^

U

r

ON ON

"a

^

PQ U U

5

I •—t

J

ffl g ffl g s^ fe

U

-H

41

-H

Q

^ a

g (/3

c 00

;ij

73 3 3 O c« O

CO

o

E

2 § .^

ON

G 3 ON


^

-H * ^ ^

-H

O

^-I .-g (^ c

si. *-• 0) (/) JH

4:^

220

Deborah A. Bronk

processes are closely coupled (Fuhrman, 1987). In Long Island Sound, DFAA supply >10% of the C and N used to fuel bacterial growth, and DFAA uptake and release rates tend to be highest near noon and lowest at night, suggesting a link to autotrophs (Fuhrman, 1987). The four amino acids measured (glutamic acid, serine, glycine, alanine) can supply 44 to 131% of the calculated bacterial N demand. In other studies, DFAA and DCAA have been shown to supply ^^50% of the bacterial N demand in estuarine and coastal systems (Keil and Kirchman 1991a, 1993; Middelboe et ai, 1995). In the subarctic Pacific and Delaware estuary, DFAAs are used preferentially over DCAAs unless DFAA concentrations are very low (Keil and Kirchman, 1991a). In 14 bioassays performed, 51 ± 45% of the bacterial N demand is met by DFAA, with 18 di 24% met by DON other than DFAA. In the Northern Sargasso Sea, protein is the dominant form of DON fueling bacterial production, supporting 20 to 65% of the estimated bacterial N demand in the surface (Keil and Kirchman, 1999). Middelboe et al (1995) also found that DFAA and DCAA sustains up to 34 and 24% of the bacterial N demand, respectively, during exponential growth. As DFAA and NHU"^ concentrations decrease during stationary phase, the importance of DCAA as both a C and an N source increases. In the Mississippi plume, rapid DFAA turnover occurs coincident with rapid NH4"^ regeneration rates, suggesting that DFAA are important substrates for bacterial NH4"'' regeneration in the plume (Cotner and Gardner, 1993). Similar findings where DFAA turnover exceeds bacterial N demand have been observed in another study in the plume (Gardner et al, 1993), in Chesapeake Bay (Fuhrman, 1990), and in the subarctic Pacific (Kirchman et al, 1989; Keil and Kirchman, 1991a). The role of DFAA as a N source for phytoplankton was reviewed in Flynn and Butler (1986) and Antia et al (1991). Though laboratory studies show that some phytoplankton can grow on DFAA, uptake of DFAA by phytoplankton is considered to be insignificant in the field; as noted above, recent research on cell-surface enzymes suggests that phytoplankton use of DFAA may be greater than previously thought (see Section IV.A.2). In a salt marsh phytoplankton conmiunity, addition of organic N, including glycine, glutamic acid, and an amino acid mixture, results in increased phytoplankton growth (Lewitus et al, 2000). The physiological response of the phytoplankton community to organic N additions, in the presence and absence of antibiotics, suggests that the stimulation caused by organic N additions results directly from uptake of the organic substrates and indirectly through bacterial decomposition. The newly recognized Archaea also appear to use DFAA. In studies in the Mediterranean Sea and the Pacific Ocean near California, ^60% of the Archaea exhibit measurable DFAA uptake at nanomolar levels (Ouvemey and Fuhrman, 2000). There is increasing recognition that the utilization of DCAA and DFAA may be affected by abiotic reactions. Glucosylation and adsorption processes appear to be

Dynamics of DON

111

important in making labile compounds more refractory. Rates of protein utilization decrease when the protein is adsorbed to submicrometer particles (Nagata and Kirchman, 1996). This is potentially a very important mechanism because the surface area of colloids in the surface ocean likely exceeds that of bacteria (Schuster et al, 1998). Accordingly, a given amino acid released from a phytoplankton cell is much more likely to come into contact with colloidal material, rendering it less biologically available, than to come into direct contact with a bacterial cell. These studies suggest that competition between abiotic adsorption onto colloids and bacterial uptake can have large implications for the cycling of DOM, particularly small labile moieties such as amino acids. An estimated ~ 11-55% of the DFAA detectable by HPLC may be adsorbed to colloidal DOM in oceanic surface waters (Schuster et al, 1998). Natural bacterial populations degraded ~92% of dissolved unprotected proteins in 72-90 h in one study (Borch and Kirchman, 1999). Protein adsorbed to or present within liposomes, designed to mimic protein that is adsorbed or trapped within particles similar to those produced by protists, however, has substantially lower degradation rates. The fecal pellets of some flagellates are believed to be similar in structure to liposomes (Nagata and Kirchman, 1992), and viral lysis can also produce liposome-like structures (Shibata et al, 1997). Reduction in the degradation rates of organics associated with liposomelike structures may explain the presence of membrane proteins in the deep ocean DOM pool (Tanoue et al, 1996; McCarthy et al, 1998). On the flip side, adsorption of DFAA can also make refractory organics more bioavailable. Adsorption of DFAA to dextran and phytoplankton-derived colloidal DOM results in approximately three times more efficient utilization of dextran or colloidal DOM by marine bacteria when compared to dextran or DOM without adsorbed DFAA (Schuster et al, 1998). 4. Humic Substances Humic substances constitute a large reservoir of organic C and N in both aquatic and terrestrial systems (Mantoura et al, 1978). Humic substances have long been recognized for their ability to chelate organometallic substances, thereby making trace metals more available to phytoplankton (Prakash, 1971; Prakash et al, 1973) and sequestering toxic heavy metals (Barber, 1973; Toledo et al, 1982). Biologically, humic substances have traditionally been considered unavailable for assimilation due to their HMW and structural complexity. More recent studies of HMW organic compounds, however, have revealed that they are not as refractory as once thought (Moran and Hodson, 1994; Amon and Benner, 1994; Gardner etal, 1996). Despite these advances, the role of marine humic substances remains unclear. It has been postulated that some phytoplankton, specifically the dinoflagellates, may be able to utilize N bound to humic substances (Carlsson and Graneli, 1993).

222

Deborah A. Bronk

Experiments in which natural humic substances, isolated from river water, are added to an assemblage of coastal phytoplankton reveal that growth and biomass formation are stimulated (Carlsson et al, 1993). The hterature suggests that the N associated with humic substances can be removed via one of three mechanisms: through microbial activity (Miiller-Wegener, 1988), via excision by phytoplankton cell-surface enzymes (Palenik and Morel, 1990a; see Section IV.A.2), or through photodegradation to LMW compounds by exposure to UV radiation (Cellar, 1986; Kieber et al, 1990; Mopper et al, 1991; see Section IV.D). 5. Other Organic Compounds Additional studies that measure uptake of other organic N compounds such as purines (Douglas, 1983), pyrimidines (Knutsen, 1972), and amines (Neilson and Larsson, 1980; Wheeler and Hellebust, 1981) show that though phytoplankton and bacteria can utilize these compounds, the uptake rates are quite low (reviewed in Antia etal,l99l). There is still a debate as to whether D-DNA is actually used as a source of N for bacteria; D-DNA is approximately 16% N and so it has the potential to be a N source. Paul et al (1988) found evidence that D-DNA is used as a source of nucleic acids for bacteria and that it is degraded to provide phosphate needed by the cell. J0rgensen et al (1993) measured uptake rates of DCAA, DFAA, and D-DNA in seawater cultures, and found that D-DNA is used primarily as a source of N. When DCAA, DFAA, and D-DNA are combined, they provide 14 to 49% of the net bacterial N uptake measured in that study. Using turnover times of unidentified HMW DON, estimated with 8^^N data, DON concentrations, and rates of primary production, Benner et al (1997) estimated that DON remineraUzation can support 30-50% of daily phytoplankton N demand in the equatorial Pacific region.

D. PHOTOCHEMICAL DECOMPOSITION AS A SINK FOR DON Recent findings in freshwater and marine systems indicate that photochemical processes can effect the release of labile N moieties from DOM (Bushaw et al, 1996). Numerous studies have shown that photochemical reactions occur when DOM from freshwater or marine environments is exposed to natural sunlight. The resulting photoproducts include carbon monoxide, carbon dioxide, various carbonyl compounds, and likely many others (see reviews by Moran and Zepp, 2000, and Mopper and Kieber, Chapter 9). Some of these photoproducts can be lost by direct transfer to the atmosphere, while others can be assimilated rapidly by natural bacterial populations (Kieber et al, 1989; Geller, 1986; Lindell et al, 1995). With respect to N, we know that substances containing organic N can play an important role in the impact of UV radiation on aquatic biogeochemical cycles (de Mora et al, 2000).

Dynamics of DON

223

To date, most of the studies of N photoproduction have focused on fresh or brackish water systems (Table VII). Studies have documented the photoproduction of NH4+, DFAA, DCAA, DPA, and NO2- (Table VII), but the process is not ubiquitous (Bertilsson et ai, 1999; Koopmans and Bronk, in press). DON and isolated humic substances can be a source of labile N when irradiated with sunlight, and wavelengths in the ultraviolet (UV) region (280400 nm) produce the N photoproducts most efficiently (Bushaw et al, 1996). Humic substances are likely important substrates for photoproduction because their aromaticity and color allow them to absorb UV light, making them more photochemically reactive than other classes of marine DOM. Furthermore, an estimated 50 to 75% of the N associated with humic substances exists as DFAA, amino sugars, and other N-rich compounds that are likely sources of the labile N forms produced photochemically (Valiela and Teal, 1979; Rice, 1982; Thurman, 1985; Stevenson, 1994). In a river and bayou in Louisiana, an estimated 9 to 20% of the TON in the photic zone was converted to NH4+ each day (Wang et al, 2000). Koopmans and Bronk (in press) measured N photoproduction from DOM isolated from surficial groundwaters. Photochemical production of NH4"^ was observed in 4 of 5 irradiated estuarine surface water samples, but in only 2 of 13 groundwater samples. In contrast, the photochemically mediated loss of NH4"^ was observed in 7 of 13 groundwater samples, likely due to incorporation into DOM. These data suggest that photochemical reactions may be a sink as well as a source of available N. In a cross-system comparison, photoproduction experiments were performed in parallel with ^^N uptake experiments (Bronk et al, unpublished data). Photochemical ammonification supplied an average of 13,13, and 7% of the NHj taken up in the Eastern Tropical North Pacific, South Atlantic Bight, and two rivers in Georgia, respectively. When photoproduction is detected, it supplies up to 38% of the DPA utihzed and up to 33% of the N02~ taken up. Photochemical ammonification is a relatively minor source of NH4'^ in all three environments with rates being 2 to 6% of biotic NH4+ regeneration rates, measured with the ^^N isotope dilution technique (Gilbert et al, 1982). In a study in Lake Maracaibo, photochemical ammonification rates are ^30% of the total near surface rates of NH4'^ regeneration (Gardner et al, 1998).

E. SINKS FOR DON: RESEARCH PRIORITIES Research on DON utihzation is poised for rapid development. Some specific areas where additional study should prove fruitful would be to address questions of the differential flow of the C and N fractions of DOM in parallel. Combining the new enzymatic approaches with dual labeled substrates (^^C, ^^N, ^^O, etc.)

224

Deborah A. Bronk Table VII

Rates of Photochemical Release from Dissolved Organic Nitrogen (DON) in Whole Water or Various DON Fractions

Substrate

Photoproduction rate (ng-atNL^h-i)

June

Isolated fiilvic acids

370 ± 10

Bushawefa/., 1996

July

Whole water

150 ± 10

Bushaw et ai, 1996

August

Isolated fulvic acids

65 ± 1 0

Bushsiw et al, 1996

Whole water

340 ± 30

Bushaw ef a/., 1996

Isolated fulvic acids

50 ± 1 5

BushsLW etai, 1996

Isolated fulvic acids

320

BushawetaL, 1996

September 1995 Whole water

0 to 220

Gardner et ai, 1998

June-Aug 1996 Whole water

ND

Bertilsson er a/., 1999

June-Aug 1996 Whole water

ND

Bertilsson e/fl/., 1999

July 1994

Whole water

ND

J0rgensen 6^ a/., 1998

< 1000 Dalton DOM < 1000 Dalton DOM

330

Wang era/., 2000

1200 to 1700

Wang era/., 2000

Location Production ofNH4+ Boreal Pond, Manitoba Boreal Pond, Manitoba Boreal Pond, Manitoba Okeefenokee Swamp, GA Satilla River, GA Oyster River, NH Lake Maracaibo, Venezuela River catchments, Sweden Groundwater, Sweden Lake Skarshult, Sweden Pearl River, LA Bayou Trepagnier, LA Bayou Trepagnier, LA Skidaway River, GA Skidaway River, GA Skidaway River, GA Satilla River, GA

Date

August 1997

Reference

January 1999

< 1000 Dalton DOM

1900

Wang et ai, 2000

August 1995

2.8 X Concentrated 2.8 X Concentrated 28 X Concentrated 2.8 X Concentrated

ND

Bushaw-Newton Moran, 1999 Bushaw-Newton Moran, 1999 Bushaw-Newton Moran, 1999 Bushaw-Newton Moran, 1999

February 1996 February 1996 October 1996

humics 7 ±4.9^ humics 60 ±3^^ humics 58 ±3^^ humics

and and and and

(Continues)

225

Dynamics of DON Table VII (Continued) Photoproduction rate (ng-atNL-^h-i)

Reference

Location

Date

Eastern Tropical North Pacific South Atlantic Bight Altamah and Savannah rivers

July 1995

Whole water

5.4 ± 4.4

Bronk et al, unpublished data

March 1999 Mar, July, Oct 1998

Whole water

35.3 ± 39.3

Whole water

10.8 ±15.1

Bronk et al, unpublished data Bronk et al, unpublished data

Substrate

Meanistd 350.0 ib 559.8^ Mean ± std 136.5 ± 139.4^^ Production of dissolved free and combined amino acids Whole water

63

J0rgensen et al. 1998

August 1995 February 1996 February 1996 October 1996 July 1995

2.8 X Concentrated humics 2.8 X Concentrated humics 28 X Concentrated humics 2.28 X Concentrated humics Whole water

ND

41 ±7.1^

Bushaw-Newton and Moran, 1999 Bushaw-Newton and Moran, 1999 Bushaw-Newton and Moran, 1999 Bushaw-Newton and Moran, 1999 Bronk et al. unpublished data

Mar, July, Oct 1998

Whole water

8.7 ± 12

Lake Skarshult, July 1994 Sweden

Production of DPA Skidaway River, GA Skidaway River, GA Skidaway River, GA Satilla River, GA Eastern Tropical North Pacific Altamah and Savannah rivers

ND

Mean ± std Production of NO2Coastal seawater, NC Albermarle sound, NC Marsh, NC Cape Fear Estuary, NC

9 ± 8.5« 6.1 ± 9 . 4

Bronk et al., unpublished data

16.2 ± 16.6

May

Isolated humics

1.4

Kithtx etal, 1999

May

Isolated humics

6.7

Kieberg^(3/., 1999

May May

Isolated humics Isolated humics

1.9 4.9

Ki&hti etal, 1999 Y^thtr etal, 1999 {Continues)

Deborah A. Bronk

226 Table VII (Continued)

Location

Date

Substrate

Photoproduction rate (ng-atNL-^h-i)

Eastern Tropical North Pacific Altamah and Savannah overs

July 1995

Whole water

4.8 ± 4.4

Bronk et al, unpublished data

Mar, July, Oct 1998

Whole water

0.3 ± 0.9

Bronk et al, unpublished data

Reference

Meanistd 3.3 ± 2.5

Note. Data are presented as mean ± standard deviation unless otherwise noted. ND: not detected. ^Standard errors. ^Including all data. ^Excluding the Bayour Trepagnier data.

will likely show that the fate of the separate elements in DOM are different trophic levels (for example, see Fig. 6). It may also show that mixotrophy is more widespread than presently recognized. Along these same lines, quantifying where the DON is going, into autotrophic versus heterotrophic biomass, is extremely important to determining how these flows are modeled. Combining tracer techniques withflowcytometric sorting is one very promising way to discriminate between autotrophic and heterotrophic uptake (Lipschultz, 1995). The increasing availability of flow cytometers and the higher sorting speeds they can reach should make this approach much more widespread in the future. Finally, the long-term goal of bringing molecular techniques to bear on issues of elemental cycling is beginning to pay off. Quantitative PCR-type approaches will continue to be refined, holding out the tantalizing possibility of estimating flux rates without the perturbations inherent in traditional incubation techniques.

V. DON TURNOVER TIMES Considering the heterogeneous nature of the DON pool, interpreting DON turnover times can be difficult. Turnover times for organic N cover a broad range from minutes for DFAA (Fuhrman, 1990) to hundreds of years for the bulk DON pool (Vidal et al, 1999; Table VIII). In the Chesapeake Bay plume, DFAAs cycle rapidly with turnover times of 0.5 to 1.0 h in spring and summer and ^ 3 h in winter (Fuhrman, 1990). When considering the bulk DON pool, Abell et al (2000) estimated turnover times, based on the surface concentrations of bioavailable TON in the mixed layer, to be 18 years when both shallow or

Dynamics of DON

227

deep isopycnal degradation estimates are used. The residence time of DON in the surface waters of the equatorial Atlantic is estimated at 2.5 years (Vidal et al, 1999). Harrison et al. (1992) estimated a maximum DON turnover time of 333 days (0.003 day~^) in the northeastern Pacific by measuring changes in DON concentrations between cruises. Considering the enormous range of turnover times, one tends to wonder whether turnover times for the bulk DON pool really tell us much. One danger in interpreting DON turnover times estimated with ^^N tracers is the convention that the shorter the turnover time, the more labile the compound. For example, Bronk and Ward (1999) found that DON turnover times, estimated with release rates measured in ^^NH4+ incubations, are shorter than those measured in incubations with ^^NOa". These data imply that DON resulting from NH4+ uptake is more labile than that resulting from NOa" uptake. In reality the compounds produced and released are likely the same in both cases because the first step after NO3" is taken up by phytoplankton is the reduction to NH4+. The lability should be the same, regardless of the substrate, because the compounds released should be the same.

VL SUMMARY Traditionally, DON has been viewed as a large refractory pool that is unimportant to microbial nutrition. Research over the past decade has transformed this view, however, and the DON pool is emerging as a dynamic component of the DOM and N cycles. It is increasingly included as a core measurement in field programs and sophisticated chemical analyses are beginning to define its structure, chemical properties, sources, and sinks. I have attempted to describe recent findings in each of these areas, which I summarize below. 1. Concentration and Composition of the DON Pool The lowest DON concentrations are generally found in the deep ocean with the highest observed in rivers (Fig. 1). DON generally accounts for the largest percentage of the TDN pool (~60%) in most systems. Though much work still needs to be done to define the global distributions of DON, the general trends emerging are that upwelling at the equator, in both the Atlantic and Pacific, fuels DON production. The DON produced is then exported to the north and south into the oligotrophic gyres. Concentrations tend to decrease near the poles, though seasonal accumulations in spring are likely, and increase near the continental margins. Vertical profiles of DON generally show a surface enrichment, and DON concentrations tend to be inversely correlated with NOs" concentrations as depth increases. Concentrations of DON and NOa" are also often inversely correlated over time in surface waters. Recent studies estimate that up to 80% of the net NOs" drawdown in a number of

228

Deborah A. Bronk Table VIII Tlimover Time Estimates of Dissolved Organic Nitrogen (DON) and Organic N Compounds Turnover time

Units

Method

DON

0.91

Years

CC

Oct-Nov 1995

DON

0.4tol3.2«

Years

cc

Oct-Nov 1995

DON

12.7 ±26.1''

Years

CC

Vidal et al, 1999

Oct-Nov 1995

DON

2.1to300''

Years

cc

Vidal et al, 1999

November 1988 October 1992

DON

40.7 ± 10.4

Days

15N

DON

11 to 62

Days

15N

July 1990, Feb 1991

Protein

0.38 to 3.42

Days

14c

Northern Sargasso Sea

July 1990

Modified 9.04 to 32.71 protein^

Days

14c

Northern Sargasso Sea

February 1991

Modified 9.04 to 32.71 protein^

Days

14c

Northern Sargasso Sea

July 1990, Feb 1991

DFAA

0.03 to 0.29

Days

3H

Central Arctic

July-Aug 1994

DFAA

-2.72

Days

^H

March 1993

DON

5.0 ± 2.4

Days

15N

September 1993 October 1992

DON

8.2 ± 2.4

Days

15N

DON

24 to 85

Days

15N

February 1991

DFAA

0.013 to 0.073'

Days

3H

Location Oceanic Northeastern Pacific Equatorial Atlantic (15S-25N) Equatorial Atlantic (15S-15N) Equatorial Atlantic (35-15S) Caribbean Sea Southern California Bight Northern Sargasso Sea

Coastal Monterey Bay Monterey Bay Southern California Bight Mississippi River plume

Date NP

Compound considered

Reference Harrison et al, 1992 Vidal et al, 1999

Bronk et al, 1994 Bronk et al, 1994 Keil and Kirchman, 1999 Keil and Kirchman, 1999 Keil and Kirchman, 1999 Keil and Kirchman, 1999 Rich era/., 1997

Bronk and Ward, 1999 Bronk and Ward, 1999 Bronk et al, 1994 Cotner and Gardner, 1993 {Continues)

Dynamics of DON

229 Table VIII {Continued)

Location

Date

Mississippi River plume Santa Rosa Sound, FL Flax Pond, NY

September 1991

Estuarine Chesapeake Bay Chesapeake Bay, mesohaline Chesapeake Bay, mesohaline Chesapeake Bay, mesohaline Chesapeake Bay, mesohaline Choptank River^ Choptank River'^ Chesapeake Bay, mesohaline Chesapeake Bay, mesohaline Chesapeake Bay, mesohaline Thames Estuary

Chesapeake Bay Chesapeake Bay

Turnover time

Units

Method

Reference

DFAA

0.02 to 0.14^

Days

3H

D-DNA

0.2 to 0.43

Days

3H

D-DNA

0.64 to 9.7

Days

3H

Cotner and Gardner, 1993 J0rgensen et al, 1993 J0rgensen et al, 1993

Days

15N

Compound considered

91.0

Bronk et al, 1994 Bronk et al. 1993a

April 1989 DON and 1990 August DON 1991

6.0 to

2.0 to 6.0

Days

15N

May 1988

DON

0.27 ± 0.23

Days

15N

Bronk et al, 1998

August 1988

DON

2.01 ±

1.13

Days

15N

Bronk et al, 1998

October 1988

DON

2.53 it: 2.54

Days

15N

Bronk et al. 1998

August 1990 August 1990 May 1988

DON

33.8

Days

15N

LMWDON

15.9

Days

15N

Urea

0.12 ± 0.03

Days

15N

Bronk et al. 1993b Bronk et al. 1993b Bronk et al, 1998

August 1988

Urea

0.33 ± 0.33

Days

15N

Bronk et al. 1998

October 1988

Urea

1.00 ± 0.30

Days

15N

Bronk et al. 1998

February 1999

Urea

4.2 to

69.0

Days

15N

1973

Urea

3.17 ± 0.63

Days

15N

1988-1997

Urea

1.10 ± 0.71

Days

15N

Middelburg and Nieuwenhuize, 2000 Lomas et al. in press Lomas et al. in press {Continues)

230

Deborah A. Bronk Table Vm (Continued)

Location

Date

Compound considered

Thames Estuary

Turnover time Units Method 0.2 to 1.9 Days

15N

3H

Middelburg and Nieuwenhuize, 2000 Furhman, 1990

Days

3H

Furhman, 1990

Days

3H

Furhman, 1990

Days

3H

Furhman, 1990

Days

3H

Furhman, 1990

1999 Hudson River plume Chesapeake Bay plume Chesapeake Bay plume Chesapeake Bay plume Chesapeake Bay plume

September 1985 February 1985 June 1985 August 1985 April 1986

glu, gly, ala glu, gly, ser, ala glu, gly, ser, ala glu, gly, ser, ala

0.060 to 0.210 0.009 to 0.090 0.017 to 0.170 0.016 to 0.240

Reference

Note. Data are presented as mean ± standard deviation. NP: not presented. ^Estimated with DON concentrations and vertical flux estimates. ^Glucosylated (i.e., aged) protein as in Keil and Kirchman (1993). ^In general, turnover times increased with salinity. ^Subestuary of Chesapeake Bay.

systems accumulates as DON. In the most general sense, a generic DON pool is shaping up to look like this: Identifiable LMW compounds such as urea, DCAA, and DFAA make up --5 to 10% of the total DON pool each. Roughly 30% of the pool is HMW (> 1 kDa). Of that HMW fraction, -20-30% is hydrolyzable amino acids with the remainder being amide in form. This leaves a substantial fraction of the pool yet to be identified 2. Sources of DON With respect to sources of DON, this review focuses on biotic water colunm processes that result in DON production from phytoplankton and N2 fixers (passive diffusion, active release, sloppy feeding, and viral lysis), bacteria (passive diffusion, release of exoenzymes, bactivory, and viral lysis), and micro- and macrozooplankton (fecal pellet dissolution and excretion; Fig. 3). Rates of DON release summarized here suggest that the magnitude of release is similar in oceanic and coastal environments but slightly higher in estuarine systems. The percentage of the rate of gross N uptake released as DON was highest in oceanic systems (—40%) and lowest in estuaries (—23%), though clearly more data are needed before these generalizations can be considered robust. 3. Sinks for DON With respect to DON sinks, this review focuses on heterotrophic uptake, autotrophic uptake, and photochemical N decomposition. Though heterotrophs have been traditionally considered the primary users of DON, there is increasing

Dynamics of DON

231

recognition that DON can be an important source of N for phytoplankton. The recent work on phytoplankton cell surface enzymes has provided a mechanism by which autotrophs can utilize the N associated with DON without developing transport mechanisms for a wide range of compounds. Much of the interest in DON uptake of late has been encouraged by a number of studies that have documented a link between increases in DON concentrations and blooms of harmful algae. Rates of DON utilization vary widely across systems and even within systems. The work summarized here suggests that the large DON pool is more bioavailable than previously thought. Work to date (much of which was done in freshwater systems with dark bioassays) suggests that 12 to 72% of the DON pool is bioavailable on the time scale of days to weeks. Three key substrates within the DON pool are urea, DCAA, and DFAA. In studies where the uptake of these substrates are compared to other N compounds, urea averages 19% of total measured N uptake with 38 and 23% contributed by DCAA and DFAA, respectively. Nitrogen photoproduction has been demonstrated in a number of environments, and it can be an important mechanism for converting DON into labile compounds available for uptake by either phytoplankton or bacteria. Photochemical anmionification has been the most studied with an average rate of 136 ng-at N L~^h~^ with some extremely high rates documented. Rates of DPA and N02~ photoproduction have tended to be lower, though only a small number of studies have been done.

ACKNOWLEDGMENTS I thank N. O. G. J0rgensen for his thought provoking review, S. Seitzinger for editorial advice, two anonymous reviewers for insightful comments, D. Karl for help with DNA/RNA calculations, and C. Carlson and D. Hansell for their patience. This work was supported by Georgia Sea Grant (NA06RG0029) and the National Science Foundation (OCE-0095940). This paper is VIMS Contribution 2400 from the Virginia Institute of Marine Science, College of William and Mary.

REFERENCES Abell, J., Emerson, S., and Renaud, R (2000). Distributions of TOP, TON, and TOC in the North Pacific subtropical gyre: Implications for nutrient supply in the surface ocean and remineralization in the upper thermocline. /. Mar. Res. 58, 203-222. Abdul-Rashid, M. (1990). Ph.D. thesis, University of Liverpool. Admiraal, W., Peletier, H. and Laane, R. P. M. (1986). Nitrogen metabolism of marine planktonic diatoms: Excretion, assimilation and cellular pools of free amino acids in seven species with different cell size. J. Exp. Mar. Biol. Ecol. 98, 241-263. Agusti, S., Satta, M. P., Mura, M. P., and Benavent, E. (1998). Dissolved esterase activity as a tracer of phytoplankton lysis: Evidence of high phytoplankton lysis rates in the northeastern Mediterranean. Limnol. Oceanogn 43, 1836-1849.

232

Deborah A Bronk

Aiken, G. R. (1988). A critical evaluation of the use of macroporous resins for the isolation of aquatic humic substances. In "Humic Substances and Their Role in the Environment" (F. H. Frimmel and R. F. Christman, Eds.), pp. 15-28. Wiley, New York. Alberts, J. J., and Takacs, M. (1999). Importance of humic substances for carbon and nitrogen transport into southeastern United States estuaries. Organ. Geochem. 30, 385-395. Alonso, M. C , Rodriguez, V., Rodriguez, J., and Borrego, J. J. (2000). Role of ciHates, flagellates and bacteriophages on the mortality of marine bacteria and on dissolved-DNA concentrations in laboratory experimental systems. /. Exp. Mar. Biol. Ecol. 244,239-252. Altabet, M. A. (1988). Variations in nitrogen isotopic composition between sinking and suspended particles: Implications for nitrogen cycling and particle transformations in the open ocean. Deep Sea Res. 35,535-554. Aluwihare, L., and Repeta, D. (1999). A comparison of the chemical characteristics of oceanic DOM and extracellular DOM produced by marine algae. Mar. Ecol. Prog. Sen 186,105-117. Aluwihare, L., Repeta, D., and Chen, R. (1997). A major biopolymeric component to dissolved organic carbon in surface seawater. Nature 387,166-169. Amon, R. M. W., and Benner, R. (1994). Rapid cycling of high-molecular-weight dissolved organic matter in the ocean. Nature 369,549-551. Amon, R., Fitznar, H. R, and Benner, R. (2001). Linkages among the bioreactivity, chemical composition, and diagenetic state of marine dissolved organic matter. Limnol. Oceanogr. 46,287-297. Antia, N. J., Harrison, R J., and Oliveira, L. (1991). Phycological reviews: The role of dissolved organic nitrogen in phytoplankton nutrition, cell biology, and ecology. Phycologia 30, 1-89. Armstrong, F. A. J., and Tibbitts, S. (1968). Photochemical combustion of organic matter in sea water for nitrogen, phosphorus, and carbon determination. J. Man biol. Assoc. UK4H, 143-152. Armstrong, R A. J., WiUiams, P. M., and Strickland, J. D. H. (1966). Photo-oxidation of organic matter in sea water by ultra-violet radiation, analytical and other appUcations. Nature 211,481^83. Axler, R. R, and Renter, J. E. (1986). A simple method for estimating the ^^N content of DOM (DO^^N) in N cycling studies. Can. J. Fish. Aquat. Sci. 43, 130-133. Baines, S. P., and Pace, M. L. (1991). The production of dissolved organic matter by phytoplankton and its importance to bacteria: Patterns across marine and freshwater systems. Limnol. Oceanogr. 36,1078-1090. Barber, R. T. (1973). Organic ligands and phytoplankton growth in nutrient-rich seawater. In "Trace Metals and Metal-Organic Interactions in Natural Waters" (P. C. Singer, Ed.), pp. 321-338. Ann Arbor Sci., Ann Arbor, MI. Bates, N. R., and Hansell, D. A. (1999). A high resolution study of surface layer hydrographic and biogeochemical properties between Chesapeake Bay and Bermuda. Man Chem. 64, 1-16. Benner, R., Biddanda, B., Black, B., and McCarthy, M. (1997). Abundance, size distribution, and stable carbon and nitrogen isotopic compositions of marine organic matter isolated by tangentialflow ultrafiltration. Mar. Chem. 57,243-263. Benner, R., Chin-Leo, G., Gardner, W, Eadie, B., and Cotner, J. (1992a). The fates and effects of riverine and shelf-derived DOM on Mississippi river plume/Gulf shelf processes. In Nutrient Enhanced Coastal Ocean Productivity, pp. 84-94. Sea Grant Program, Texas A&M University, Galveston, TX. Benner, R., Pakulski, J. D., McCarthy, M., Hedges, J. I., and Hatcher, R G. (1992b). Bulk chemical characteristics of dissolved organic matter in the ocean. Science 255,1561-1564. Benner, R. H. (2002). Composition and Reactivity. In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. Hansell and C.A. Carlson, Eds.), pp. 58-90. Academic Press, San Diego. Berg, G. M., Gilbert, P. M., Lomas, M. W, and Burford, M. A. (1997). Organic nitrogen uptake and growth by the chrysophyte Aureococcus anophagejferens during a brown tide event. Mar Biol. 129, 377-387. Berges, J. A., and Falkowski, P. G. (1996). CeU-associated proteolytic enzymes from marine phytoplankton. J. Phycol. 32, 566-574.

Dynamics of DON

233

Berman, T., Bechemin, C , and Maestrini, S. Y. (1999). Release of ammonium and urea from dissolved organic nitrogen in aquatic ecosystems. Aquat. Microb. Ecol 16, 295-302. Berman, T., and Chava, S. (1999). Algal growth on organic compounds as nitrogen sources. J. Plankton Res. 21,1391-1421. Berman, T., Chava, S., Kaplan, B., and Wynne, D. (1991). Dissolved organic substrates as phosphorus and nitrogen sources for axenic batch cultures of freshwater green algae. Phycologia 30, 339345. Bertilsson, S., Stepanauskas, R., Cuadros-Hansson, R., GraneH, W., Wikner, J., and Tranvik, L. (1999). Photochemically induced changes in bioavailable carbon and nitrogen pools in a boreal watershed. Aquat. Microb. Ecol. 19,47-56. Biddanada, B., and Benner, R. (1997). Carbon, nitrogen and carbohydrate fluxes during the production of particulate and dissolved organic matter by marine plankton. Limnol Oceanogr. 42,506-518. Billen, G. (1991). Protein degradation in aquatic environments. In "Microbial Enzymes in Aquatic Environments" (R. J. Chrost, Ed.), pp. 123-143. Springer-Verlag, New York. Bj0msen, P. K. (1988). Phytoplankton exudation of organic matter: Why do healthy cells do it? Limnol. Oceanogr. 33, 151-155. Borch, N. H., and Kirchman, D. L. (1999). Protection of protein from bacterial degradation by submicron particles. Aquat. Microb. Ecol. 16, 265-272. Branch, G. M., and Griffiths, C. L. (1988). The Benguela ecosystem part V. The coastal zone. Oceanogr. Mar. Biol. Annu. Rev. 26, 395-486. Bratbak, G., Heldal, M., Naess, A., and Roeggen, T. (1993). Viral impact on microbial communities. In "Trends in Microbial Ecology" (R. Guerrero and C. Pedros-AUo, Eds.), pp. 299-302. Spanish Society for Microbiology, Barcelona, Spain. Bratbak, G., Jacobsen, A., and Heldal, M. (1998). Viral lysis of Phaeocystis pouchetii and bacterial secondary production. A^war. Microb. Ecol. 16,11-16. Bratbak, G., Thingstad, R, and Heldal, M. (1994). Viruses and the microbial loop. Microb. Ecol. 28, 209-221. Bratbak, G., and Thingstad, T. F. (1985). Phytoplankton-bacteria interactions: An apparent paradox? Analysis of a model system with both competition and conmiensalism. Mar Ecol. Prog. Sen 25, 23-30. Bricelj, V. M., and Lonsdale, D. J. (1997). Aureococcus anophagefferens: Causes and ecological consequences of brown tides in U.S. mid-Atlantic coastal waters. Limnol. Oceanogr 42,1023-1038. Bronk, D. A. (1999). Rates of NH4+ uptake, intracellular transformation, and dissolved organic nitrogen release in two clones of marine Synechococcus spp. /. Plankton Res. 21,1337-1353. Bronk, D. A., and Gilbert, P. M. (1991). A ^^N tracer method for the measurement of dissolved organic nitrogen release by phytoplankton. Mar Ecol. Prog. Sen 11,171-182. Bronk, D. A., and Gilbert, P. M. (1993a). Application of a ^^N tracer method to the study of dissolved organic nitrogen uptake during spring and summer in Chesapeake Bay. Mar Biol. 115,501-508. Bronk, D. A., and Gilbert, P. M. (1993b). Contrasting patterns of dissolved organic nitrogen release by two size fractions of estuarine plankton during a period of rapid NH4+ consumption and NO^ production. Mar Ecol. Prog. Sen 96, 291-299. Bronk, D. A., Gilbert, P. M., and Ward, B. B. (1994). Nitrogen uptake, dissolved organic nitrogen release, and new production. Science 265,1843-1846. Bronk, D. A., Gilbert, P. M., Malone, T. C , Banahan, S., and Sahlsten, E. (1998). Inorganic and organic nitrogen cycling in Chesapeake Bay: Autotrophic versus heterotrophic processes and relationships to carbon flux. Aquat. Microb. Ecol 15, 177-189. Bronk, D. A., Lomas, M., Gilbert, R M., Schukert, K. J., and Sanderson, M. P (2000). Total dissolved nitrogen analysis: Comparisons between the persulfate, UV and high temperature oxidation method. Mar Chem. 69,163-178. Bronk, D. A., and Ward, B. B. (1999). Gross and net nitrogen uptake and DON release in the euphotic zone of Monterey Bay, California. Limnol. Oceanogr 44, 573-585.

234

Deborah A. Bronk

Bronk, D. A., and Ward, B. B. (2000). Magnitude of DON release relative to gross nitrogen uptake in marine systems. Limnol Oceanogr. 45,1879-1883. Brown, M. R. (1991). The amino acid and sugar composition of 16 species of microalgae used in mariculture. J. Exp. Mar. Biol. Ecol. 145,79-99. Burdige, D. J. (2002). Sediment Pore Waters. In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. Hansell and C.A. Carlson, Eds.), pp. 611-663. Academic Press, San Diego. Bury, S. J., Owens, N. J. P., and Preston, T. (1995). ^'^C and ^^N uptake by phytoplankton in the marginal ice zone of the Bellingshausen Sea. Deep-Sea Res. II42,1225-1252. Bushaw, K. L., Zepp, R. G., Tarr, M. A., Schulz-Jander, D., Bourbonniere, R. A., Hodson, R., Miller, W. L., Bronk, D. A., and Moran, M. A. (1996). Photochemical release of biologically labile nitrogen from dissolved organic matter. Nature 381,404-407. Bushaw-Newton, K. L., and Moran, M. A. (1999). Photochemical formation of biologically available nitrogen from dissolved humic substances in coastal marine systems. Aquat. Microb. Ecol 18, 285-292. Butler, E. I., Knox, S., and Liddicoat, M. I. (1979). The relationship between inorganic and organic nutrients in seawater. J. Mar. Biol. Assoc. UK 59, 239-250. Capone, D. G. (2000). The marine nitrogen cycle. In "Microbial Ecology of the Oceans" (D. L. Kirchman, Ed.), pp. 455^94. Wiley, New York. Capone, D., Ferrier, M., and Carpenter, E. (1994). Amino acid cycUng in colonies of the planktonic marine cyanobacterium Trichodesmium thiehautii. Appl Environ. Microbiol. 60, 3989-3995. Capone, D., Zehr, J., Paerl, H., Bergman, B., and Carpenter, E. (1997). Trichodesmium, a globally significant marine cyanobacterium. Science 276,1221-1229. Carlson, C. A. (2002). Production and removal processes. In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. Hansell and C.A. Carlson, Eds.), pp. 91-151. Academic Press, San Diego. Carlson, C. A., and Ducklow, H. W. (1995). Dissolved organic carbon in the upper ocean of the central equatorial Pacific Ocean, 1992: Daily and fine scale vertical variations. Deep-Sea Res. II 42, 639-656. Carlson, C. A., Ducklow, H. W, and Michaels, A. F. (1994). Annual flux of dissolved organic carbon from the euphotic zone in the northwestern Sargasso Sea. Nature 371,405-408. Carlson, C. A., Hansell, D. A., Peltzer, E. T., and Smith, Jr., W O. (2000). Stocks and dynamics of dissolved and particulate organic matter in the Southern Ross Sea, Antarctica. Deep Sea Res. II42, 3201-3225. Carlsson, P., Edling, H., and Bechemin, C. (1998). Interactions between a marine dinoflagellate (Alexandrium catenella) and a bacterial community utilizing riverine humic substances. Aquat. Microb. Ecol. 16,65-80. Carlsson, P., and GraneU, E. (1993). AvailabiHty of humic bound nitrogen for coastal phytoplankton. Estuarine Coastal Shelf Sci. 36,433^^7. Carlsson, P., GraneH, E., Tester, P., and Boni, L. (1995). Influences of riverine humic substances on bacteria, protozoa, phytoplankton, and copepods in a coastal plankton community. Mar. Ecol. Prog. Sen 127,213-221. Carlsson, P., Segatto, A. Z., and Graneli, E. (1993). Nitrogen bound to humic matter of terrestrial origin—A nitrogen pool for coastal phytoplankton? Mar Ecol. Prog. Ser. 97, 105-116. Caron, D. A. (2000). Symbiosis and mixotrophy among pelagic microorganisms. In "Microbial Ecology of the Oceans" (D. L. Kirchman, Ed.), pp. 495-523. Wiley-Liss, New York. Cherrier, J., Bauer, J. E., and Druffel, E. R. M. (1996). Utilization and mmover of labile dissolved organic matter by bacterial heterotrophs in eastern North Pacific surface waters. Mar Ecol. Prog. Ser 139,267-279. Cho, B., Park, M., Shim, J., and Azam, F. (1996). Significance of bacteria in urea dynamics in coastal surface waters. Mar Ecol. Prog. Ser 142, 19-26.

Dynamics of DON

235

Cho, B.C., and Azam, F. (1995). Urea decomposition by bacteria in the Southern California Bight and its implications for the mesopelagic nitrogen cycle. Mar. Ecol. Prog. Sen 122, 21-26. Chrost, R. J. (1991). Environmental control of the synthesis and activity of aquatic microbial ectoenzymes. In "Microbial Enzymes in Aquatic Environments" (R. J. Chrost, Ed.), pp. 25-59. Springer, New York. Cochlan, W. R, and Harrison, P. J. (1991). Kinetics of nitrogen (nitrate, ammonium, and urea) uptake by the picofiagellaiQ Micromonas pusilla (Prasinophyceae). J. Exp. Mar Biol. Ecol. 153,129-141. CoUos, Y. (1986). Time-lag algal growth dynamics: Biological constraints on primary production in aquatic environments. Mar Ecol. Prog. Ser 33, 193-206. CoUos, Y, Descolas-Gros, C , and Momet, E (1996). Diel variation in dissolved organic nitrogen in a coastal pond: Relationships with carbon and nitrogen metabolism. /. Exp. Mar Biol. Ecol. 202, 177-189. CoUos, Y, Dohler, G., and Biermann, I. (1992). Production of dissolved organic nitrogen during uptake of nitrate by Synedra planctonica: Implications for estimating new production in the oceans. /. Plankton Res. 14,1025-1029. Conover, R. J., and Gustavson, K. R. (1999). Sources of urea in arctic seas: zooplankton metabolism. Mar Ecol. Prog. Ser 179, 41-54. Conover, R. J., Mumm, N., Bruecker, P., and MacKenzie, S. (1999). Sources of urea in arctic seas: Seasonal fast ice? Mar Ecol. Prog. Ser 179, 55-69. Cornell, S., Rendell, A., and Jickells, T. (1995). Atmospheric inputs of dissolved organic nitrogen to the oceans. Nature 376,243-246. Cotner, J, B., and Gardner, W. S. (1993). Heterotrophic bacterial mediation of ammonium and dissolved free amino acid fluxes in the Mississippi River plume. Mar Ecol. Prog. Ser 93, 75-87. Cowan, S. W., Schirmer, T., Rummel, G., Steiert, M., Chosh, R., Pauptit, R. A., Jansonius, J. N., and Rosenbusch, J. P. (1992). Crystal structures explain functional properties of two E. coli porins. Nature 35S, 121-133. Cox, R. D. (1980). Determination of nitrate and nitrite at the parts per billion level by chemiluminescence. Anal. Chem. 52, 332-335. Cresser, M. S. (1977). Nitrate determination by reduction to ammonia and gas-phase ultraviolet adsorption spectrometry. Analyst 102,99-103. Dagg, M. J. (1974). Loss of prey body contents during feeding by an aquatic predator. Ecology 55, 903-906. Daly, K. L., Wallace, D. W. R., Smith, J., W. O., Skoog, A., Lara, R., Gosselin, M., Falck, E., and Yager, P. L. (1999). Non-Redfield carbon and nitrogen cycling in the Arctic: Effects of ecosystem structure and dynamics. J. Geophys. Res. 104, 3185-3199. De Mora, S., Demers, S., and Vemet, M. (2000). "The Effects of UV Radiation in the Marine Environment." Cambridge University Press, Cambridge. Derenne, S., Largeau, C , and Taulelle, F. (1992). Occurrence of non-hydrolysable amides in the macromolecular constituents of Scenedesmus quadricauda cell wall as revealed by ^^N NMR: Origin of n-alkylnitriles in pyprolysates of ultralaminae-containing kerogens. Geochim. Cosmochim. Acta 57,851-857. Diaz, F., and Raimbault, P. (2000). Nitrogen regeneration and dissolved organic nitrogen release during spring in a NW Mediterranean coastal zone (Gulf of Lions): Implications for the estimation of new production. Mar Ecol. Prog. Ser 197, 51-65. Douglas, A. E. (1983). Uric acid utilization in Platymonas convolutae and symbiotic Convoluta roscoffensis. J. Mar Biol. Assoc. UK 63,435-447. Dzurica, S., Lee, C , and Cosper, E. M. (1989). Role of environmental variables, specifically organic compounds and micronutrients, in the growth of the chrysophyte Aureococcus anaphagefferens. In "Novel Phytoplankton Blooms" (E. M. Cosper, V. M. Bricelj, and E. J. Carpenter, Eds.), pp. 229-252. Spriner-Veriag, Beriin.

236

Deborah A. Bronk

Ferrier-Pages, C , Kamer, M., and Rassoulzadegan, F. (1998). Release of dissolved amino acids by flagellates and ciliates grazing on bacteria. Oceanol Acta 21,485-494. Feuerstein, T., Ostrom, R, and Ostrom, N. (1997). Isotope biogeochemistry if dissolved organic nitrogen: A new technique and application. Org. Geochem. 27, 363-370. Flynn, K. J. (1990). Composition of intracellular and extracellular pools of amino acids, and amino acid utilization of microalgae of different sizes. J. Exp. Mar. Biol. Ecol. 139,151-166. Flynn, K. J., and Berry, L. S. (1999). The loss of organic nitrogen during marine primary production may be significantly overestimated when using ^^N substrates. Proc. R. Soc. London B 266, 641647. Flynn, K. J., and Butler, I. (1986). Nitrogen sources for the growth of microalgae: role of dissolved free amino acids. Mar Ecol. Prog. Sen 34, 281-304. Flynn, K. J., and Davidson, K. (1993). Predator-prey interactions between Isochrysis galbana and Oxyrrhis marina. 11. Prelease of non-protein amines and faeces during predation of Isochrysis. J. Plankton Res. 15, 893-905. Fogg, G. E. (1966). The extracellular products of algae. Oceanogr. Mar Biol. Ann. Rev. 4,195-212. Fogg, G. E. (1983). The ecological significance of algal extracellular products of phytoplankton photosynthesis. Bot. Mar 26, 3-14. Fu, T, and Pocklington, R. (1983). Quantitative adsorption of organic matter from seawater on soHd mattrices. Mar Chem. 13, 255-264. Fuhrman, J. (1987). Close coupling between release and uptake of dissolved free amino acids in seawater studied by an isotope dilution approach. Mar Ecol. Prog. Ser 37,45-52. Fuhrman, J. (1990). Dissolved free amino acid cycling in an estuarine outflow plume. Mar Ecol. Prog. Ser 66, 197-203. Fuhrman, J. (1992). Bacterioplankton roles in cycling of organic matter: The microbial loop. In "Primary Productivity and Biogeochemical Cycles in the Sea" (P. G. Falkowski and A. D. Woodhead, Eds.), pp. 361-381. Plenum Press, New York. Fuhrman, J. A. (1999). Marine viruses and their biogeochemical and ecological effects. Nature 399, 541-548. Fuhrman, J. A. (2000). Impact of viruses on bacterial processes. In "Microbial Ecology of the Oceans" (D. L. Kirchman, Ed.), pp. 327-350. Wiley-Liss, New York. Fuhrman, J. A., and Noble. (1995). Viruses and protists cause similar bacterial mortality in coastal seawater. Limnol. Oceanogr 40, 1236-1242. Gagosian, R. B., and Lee, C. (1981). Processes controlling the distribution of biogenic organic compounds in seawater. In "Marine Organic Chemistry: Evolution, Composition, Interactions and Chemistry of Organic Matter in Seawater" (E. K. Duursma and R. Dawson, Eds.), pp. 91-124. Elsevier, Amsterdam. Galloway, J. N., Schlesinger, W. H., Levy, H., Michaels, A. F , and Schnoor, J. L. (1995). Large changes in oceanic nutrient inventories from glacial to interglacial periods. Nature 376,755-758. Gardner, W., Benner, R., Amon, R., Cotner, J., Cavaletto, J., and Johnson, J. (1996). Effects of high molecular weight dissolved organic matter on the nitrogen dynamics on the Mississippi River plume. Mar Ecol. Prog. Ser 133, 287-297. Gardner, W. S., Cavaletto, J. F , Bootsma, H. A., Lavrentyev, P. J., and Tanvone, F. (1998). Nitrogen cycling rates and light effects in tropical Lake Maracaibo, Venezuela. Limnol. Oceanogr 43,18141825. Gardner, W. S., Cotner, J. B., and Herche, L. R. (1993). Chromatographic measurement of nitrogen mineralization rates in marine coastal waters with ^^N. Mar Ecol. Prog. Ser 93, 65-73. Gardner, W. S., and Stephens, J. A. (1978). Stability and composition of terrestrially derived dissolved organic nitrogen in continental shelf waters. Mar Chem. 6, 335-342. Garside, C. (1982). A chemiluminescent technique for the determination of nanomolar concentrations of nitrate and nitrite in seawater. Mar Chem. 11, 159-167.

Dynamics of DON

237

Geller, A. (1986). Comparison of mechanisms enhancing biodegradabiUty of refractory lake water constituents. Limnol. Oceanogr. 31, 755-764. Gibb, S. W., Mantoura, R. F. C , Liss, P. S., and Barlow, R. G. (1999). Distributions and biogeochemistries of methylamines and ammonium in the Arabian Sea. Deep-Sea Res. II46, 593-615. Gilbert, R M., and Bronk, D. A. (1994). Release of dissolved organic nitrogen by marine diazotrophic cyanobacteria Trichodesmium spp. Appl. Environ. Microbiol. 36, 3996-4000. Glibert, R M., Garside, C , Fuhrman, J. A., and Roman, M. R. (1991). Time-dependent coupling of inorganic and organic nitrogen uptake and regeneration in the plume of the Chesapeake Bay estuary and its regulation by large heterotrophs. Limnol. Oceanogr. 36, 895-909. Glibert, R M., Lipschultz, R, McCarthy, J. J., and Altabet, M. A. (1982). Isotope dilution models of uptake and remineralization of anmionium by marine plankton. Limnol. Oceanogr. 27,639-650. Glibert, R M., and Terlizzi, D. E. (1999). Cooccurrence of elevated urea levels and dinoflagellate blooms in temperate estuarine aquaculture ponds. Appl. Environ. Microbiol. 65, 5594-5596. Gobler, C. J., Hutchins, D. A., Fisher, N. S., Cosper, E. M., and Saiiudo-Wilhelmy, S. A. (1997). Release and bioavailability of C, N, P, Se, and Fe following viral lysis of a marine chrysophyte. Limnol. Oceanogr 42,1492-1504. Gobler, C. J., and Safiudo-Wilhelmy, S. A. (2001). Effects of organic carbon, organic nitrogen, inorganic nutrients, and iron additions on the growth of phytoplankton and bacteria during a brown tide bloom. Mar Ecol. Prog. Ser 209,19-34. Gordeev, V. V., Martin, J. M., Sidorov, I. S., and Sidorova, M. V. (1996). A reassessment of the Eurasian River input of water sediment, major elements, and nutrients to the Arctic Ocean. Am. J. Set 296, 664-691. Hammer, K. D., and Kattner, G. (1986). Dissolved free amino acids in the marine environment: a carbon to nitrogen shift during diatom blooms. Mar Ecol. Prog. Ser 31, 35-45. Hansen, D., and Waterhouse, T. (1997). Controls on the distributions of organic carbon and nitrogen in the eastern Pacific Ocean. Deep-Sea Res. 44, 843-857. Hansen, D. A. (1993). Results and observations from the measurement of DOC and DON in seawater using a high-temperature catalytic oxidation technique. Mar Chem. 41,195-202. Hansell, D. A. (2002). DOC in the Global Ocean Carbon Cycle. In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. Hansell and C.A. Carlson, Eds.), pp. 685-715. Academic Press, San Diego. Hansell, D. A., and Carlson, C. A. (2001). Biogeochemistry of total organic carbon and nitrogen in the Sargasso Sea: Control by convective overturn. Deep Sea Res. II, 48,1649-1667. Hansell, D. A., and Feely, R. A. (2000). Atmospheric intertropical convergence impacts surface ocean carbon and nitrogen biogeochemistry in the western tropical Pacific. Geophys. Res. Lett. 27,10131016. Hansell, D. A., WiUiams, R M., and Ward, B. B. (1993). Measurements of DOC and DON in the Southern California Bight using oxidation by high temperature combustion. Deep-Sea Res. 40, 219-234. Hansell, P. A., and Goering, J. J. (1989). A method for estimating uptake and production rates for urea in seawater using ^^C urea and ^^N urea. Can. J. Fish. Aquat. Sci. 46,198-202. Harrison, W. G., Harris, L. R., Karl, D. M., Knauer, G. A., and Redalje, D. G. (1992). Nitrogen dynamics at the VERTEX time-series site. Deep-Sea Res. 39,1535-1552. Harvey, G. R., and Boran, D. A. (1985). Geochemistry of humic substances in seawater. In "Humic Substances in Soil Sediment and Water" (G. R. Aiken, D. M. McKnight, R. L. Wershaw, and R MacCarthy, Eds.), pp. 223-247. Wiley, New York. Harvey, G. R., Boran, D. A., and Chesal, L. A. (1983). The structure of marine fulvic and humic acids. Mar Chem. 12,119-132. Hasegawa, T, Koike, I., and Mukai, H. (2000a). Release of dissolved organic nitrogen by sizefractionated natural planktonic assemblages in coastal waters. Mar Ecol. Prog. Ser 198, A3-A9.

238

Deborah A. Bronk

Hasegawa, T., Koike, I., and Mukai, H. (2000b). Estimation of dissolved organic nitrogen release by micrograzers in natural plankton assemblages. Plankton Biol. Ecol 47, 23-30. Hasagawa, T., Koike, I., and Mukai, H. (2000c). Dissolved organic nitrogen dynamics in coastal waters and the effect of copepods. /. Exp. Mar. Biol. Ecol. 244, 219-238. Hatch, M. J., Dillon, J. A., and Smith, H. B. (1957). Preparation and use of snake-cage poly electrolytes. Ind. Eng. Chem. 49,1812-1818. Hatcher, P. G., and Spiker, E. C. (1988). Selective degradation of plant biomolecules. In "Humic Substances and Their Role in the Environment" (F. H. Frimmel and R. F. Christman, Eds.), pp. 59-74. Wiley, New York. Hedges, J. I. (1988). Polymerization of humic substances in natural environments. In "Humic Substances and Their Role in the Environment" (F. H. Frimmel and R. F. Christman, Eds.), pp. 45-58. Wiley, New York. Hedges, J. I., and Hare, P. E. (1987). Amino acid adsorption by clay minerals in distilled water. Geochim. Cosmochim. Acta 51,255-259. Hessen, D. O., and Tranvik, L. J. (1998). "Aquatic Humic Substance: Ecology and Biogeochemistry." Springer, Berlin. Hoffman, M., and Decho, A. W. (2000). proteolytic enzymes in the marine bacterium Pseudoalteromonas atlantica: Post-secretional activation and effects of environmental conditions. Aquat. Microb. Ecol 23, 29-39. HoUibaugh, J. T., and Azam, F. (1983). Microbial degradation of dissolved proteins in seawater. Limnol. OceanogA: 28, 1104-1116. Hooper, J. K., and Hughes, M. J. (1992). Purification and characterization of a membrane-bound protease from Chlamydomonas reinhardtii. Plant Physiol. 99,932-937. Hopkinson, C , Buffam, I., Hobbie, J., Vallino, J., Perdue, M., Eversmeyer, B., Prahl, F., Covert, J., Hodson, R., Moran, M. A., Smith, E., Baross, J., Crump, B., Findlay, S., and Foreman, K. (1998). Terrestrial inputs of organic matter to coastal ecosystems: An intercomparison of chemical characteristics and bioavailability. Biogeochemistry. 43, 211-234. Hopkinson, C , Fry, B., and Nolin, A. (1997). Stoichiometry of dissolved organic matter dynamics on the continental shelf of the northeastern USA. Cont. Shelf Res. 17,473^89. Hoppe, H. G. (1983). Significance of exoenzymatic activities in the ecology of brackish water: Measurements by means of methylumbelliferyl-substrates. Mar. Ecol. Prog. Sen 11,299-308. Hu, S. H., and Smith, W O. (1998). The effects of irradiance on nitrate uptake and dissolved organic nitrogen release by phytoplankton in the Ross Sea. Cont. Shelf Res. 18,971-990. Hubberten, U., Lara, R. J., and Kattner, G. (1995). Refractory organic compounds in polar waters: Relationship between humic substances and amino acids in the Arctic and Antarctic. J. Mar. Res. 53, 137-149. letswaart, T, Schneider, P. J., and Prins, R. A. (1994). Utilization of organic nitrogen sources by two phytoplankton species and a bacterial isolate in pure and mixed cultures. Appl. Environ. Microbiol. 60,1554-1560. Ishiwatari, R. (1992). Macromolecular material (humic substances) in the water colunm and sediments. Mar. Chem. 2,% l5\-\66. Jackson, G. A., and WiUiams, P. W (1985). Importance of dissolved organic nitrogen and phosphorus to biological nutrient cychng. Deep-Sea Res. 32,223-235. Jenkinson, I. R., and Biddanda, B. A. (1995). Bulk-phase viscoelastic properties of seawater: Relationship with plankton components. J. Plankton Res. 17, 2251-2274. John, E. H., and Rynn, K. J. (1999). Amino acid uptake by the toxic dinoflagellate Alexandrium fundyense. Mar Biol. 133, 11-19. Jones, K., and Stewart, W D. P. (1969). Nitrogen turnover in marine and brackish habitats. III. The production of extracellular nitrogen by Calothrix scopulorum. J. Mar Biol. Assoc. UK 49,475^88. J0rgensen, N. O. G., Kroer, N., and Coffin, R. B. (1994). UtiUzation of dissolved nitrogen by heterotrophic bacterioplankton: Effect of substrate C/N ratio. Appl. Env. Microbiol. 60,4124-4133.

Dynamics of DON

239

J0rgensen, N. O. G., Kroer, N., Coffin, R. B., Yang, X.-H., and Lee, C. (1993). Dissolved free amino acids, combined amino acids, and DNA as sources of carbon and nitrogen to marine bacteria. Mar. Ecol Prog. Ser. 98,135-148. J0rgensen, N. O. G., Tranvik, L. J., and Berg, G. M. (1999). Occurrence and bacterial cycling of dissolved nitrogen in the Gulf of Riga, the Baltic Sea. Mar Ecol. Prog. Ser 191, 1-18. J0rgensen, N. O. G., Tranvik, L., Edling, H., Graneli, W., and Lindell, M. (1998). Effects of sunlight on occurrence and bacterial turnover of specific carbon and nitrogen compounds in lake water. FEMS Microbiol. Ecol. 25, 217-227. Jumars, R A., Penry, D. L., Baross, J. A., Perry, M. J., and Frost, B. W. (1989). Closing the microbial loop: Dissolved carbon pathway to heterotrophic bacteria from incomplete ingestion, digestion, and absorption in animals. Deep-Sea Res. 36,483-495. Kahler, P., Bj0msen, P. K., Lochte, K., and Antia, A. (1997). Dissolved organic matter and its utiHzation by bacteria during spring in the Southern Ocean. Deep-Sea Res. II44, 341-353. Kahler, P., and Koeve, W. (2001). Marine dissolved organic matter: Can its C:N ratio explain carbon overconsumption. Deep Sea Res. II48,49-62. Karl, D. M., and Bailiff, M. D. (1989). The measurement and distribution of dissolved nucleic acids in aquatic environments. Limnol. Oceanogr 34, 543-558. Karl, D. M., and Bjorkman, K. M. (2002). Dynamics of Phosphorus. In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. Hansell and C.A. Carlson, Eds.), pp. 249-366. Academic Press, San Diego. Keil, R. G., and Kirchman, D. L. (1991a). Contribution of dissolved free amino acids and ammonium to the nitrogen requirements of heterotrophic bacterioplankton. Mar Ecol. Prog. Ser 73,1-10. Keil, R. G., and Kirchman, D. L. (1991b). Dissolved combined amino acids in marine waters as determined by a vapor-phase hydrolysis method. Mar Chem. 33, 243-259. Keil, R. G., and Kirchman, D. L. (1993). Dissolved combined amino acids: Chemical form and utiHzation by marine bacteria. Limnol. Oceanogr 38,1256-1270. Keil, R. G., and Kirchman, D. L. (1994). Abiotic transformations of labile protein to refractory protein in sea water. Mar Chem. 45,187-196. Keil, R. G., and Kirchman, D. L. (1999). UtiUzation of dissolved protein and amino acids in the northern Sargasso Sea. Aquat. Microb. Ecol. 18, 293-800. Karl, D. M., Christian, J. R., Dore, J. E., and LeteUer, R. M. (1996). Microbiological oceanography in the region west of the Antarctic peninsula: Microbial dynamics, nitrogen cycle, and carbon flux. In "Foundations for Ecological Research West of the Antarctic Peninsula, Antarctic Research Series" (R. M. Ross, E. E. Hofmann, and L. B. Quetin, Eds.), Vol. 70, pp. 303-332. American Geophysical Union, Washington, DC. Karl, D. M., Letelier, R., Hebel, D. V., Bird, D. F , and Winn, C. D. (1992). Trichodesmium blooms and new nitrogen in the North Pacific Gyre. In "Marine Pelagic Cyanobacteria: Trichodesmium and Other Diazotrophs" (E. J. Carpenter, D. G. Capone, and J. G. Renter, Eds.), pp. 219-237. Kluwer Academic, Dordrecht, The Netherlands. Karl, D., Letelier, R., Tupas, L., Dore, J., Christian, J., and Hebel, D. (1997). The role of nitrogen fixation in biogeochemical cycUng in the subtropical North Pacific Ocean. Nature 388,533-538. Kemer, M., and Spitzy, A. (2001). Nitrate regeneration coupled to degradation of different size fractions of DON by the picoplankton in the Elbe Estuary. Microb. Ecol. 41, 69-81. Kieber, R. J., Li, A., and Seaton, P. J. (1999). Production of nitrite from the photodegradation of dissolved organic matter in natural waters. Environ. Sci. Technol 33,993-998. Kieber, D. J., McDaniel, J., and Mopper, K. (1989). Photochemical source of biological substrates in sea water: Implications for carbon cycHng. Nature 341,637-639. Kieber, R. J., Zhou, X., and Mopper, K. (1990). Formation of carbonyl compounds from UV-induced photodegradation of humic substances in natural waters: Fate of riverine carbon in the sea. Limnol. Oceanogr 35, 1503-1515. Kirchman, D. L., Ducklow, H. W, McCarthy, J. J., and Garside, C. (1994). Biomass and nitrogen

240

Deborah A. Bronk

uptake by heterotrophic bacteria during the spring bloom in the North Atlantic Ocean. Deep-Sea Res. II41, S19-S95. Kirchman, D. L., Keil, R. G., and Wheeler, P. A. (1989). The effect of amino acids on anmionium utilization and regeneration by heterotrophic bacteria in the subarctic Pacific. Deep-Sea Res. 36, 1763-1776. Kirchman, D. L., Lancelot, C , Fasham, M., Legendre, L., Radach, G., and Scott, M. (1993). Dissolved organic matter in biogeochemical models of the ocean. In "Towards a Model of Ocean Biogeochemical Processes," Vol. 40, pp. 209-225. Springer-Verlag, Berlin. Knutsen, G. (1972). Uptake of uracil by synchronous Chlorellafusca. Phys. Planet. 27, 300-309. Koike, I., and Tupas, L. (1993). Total dissolved nitrogen in the Northern North Pacific assessed by a high-temperature combustion method. Mar. Chem. 41, 209-214. Kokkinakis, S. A., and Wheeler, P. A. (1988). Uptake of ammonium and urea in the northeast Pacific: Comparison between netplankton and nanoplankton. Mar. Ecol. Prog. Ser. 43,113-124. Koopmans, D. J., and Bronk, D. A. (in press). Photochemical production of inorganic nitrogen from dissolved organic nitrogen in waters of two estuaries and adjacent surficial groundwaters. Aquat. Microbial. Ecol. Kroer, N., Jorgensen, N. O. G., and Coffin, R. B. (1994). Utilization of dissolved nitrogen by heterotrophic bacterioplankton: A comparison of three ecosystems. Appl. Env. Microbiol. 60, 41164123. Lampert, W (1978). Release of dissolved organic carbon by grazing zooplankton. Limnol. Oceanogn 23,831-834. Lampitt, R. S., Noji, T., and von Bodungen, B. (1990). What happens to zooplankton faecal pellets? ImpHcations for material flux. Mar. Biol. 104, 15-23. Lara, C , Rodriguez, R., and Guerrero, M. G. (1993). Sodium-dependent nitrate transport and energetics of cyanobacteria. J. Phycol 29, 389-395. Lara, R. J., Hubberten, U., Thomas, D. N., Baumann, M. E. M., and Kattner, G. (1997). Dissolved organic matter studies in enclosed systems: Application of hydrophobic fractionation for the assessment of organic nitrogen dynamics. / Mar. Syst. 13,155-161. LaRoche, J., Nuzzi, R., Waters, R., Wyman, K., Falkowski, P G., and Wallace, D. W. R. (1997). Brown Tide blooms in Long Island's coastal waters linked to interannual variability in groundwater flow. Global Change Biol. 3, 397-410. Larsson, U., and Hagstrom, A. (1979). Phytoplankton exudate release as an energy source for the growth of pelagic bacteria. Mar Biol. 52, 199-206. Lee, C , and Bada, J. L. (1977). Dissolved amino acids in the equatorial Pacific, the Sargasso Sea, and Biscayne Bay. Limnol. Oceanogn 11, 502-510. Letelier, R. M., and Karl, D. M. (1996). Role of Trichodesmium spp. in the productivity of the subtropical North Pacific Ocean. Mar Ecol. Prog. Ser 133, 263-273. Lewitus, A. J., Koepfler, E. T, and Pigg, R. J. (2000). Use of dissolved organic nitrogen by a salt marsh phytoplankton bloom community. Arch. Hydrobiol. Spec. Issues Adv. Limnol. 55,441-456. Lewitus, A. J., WiUis, B. M., Hayes, K. C , Burkholder, J. M., Glasgow, J., H. B., Gilbert, P M., and Burke, M. K. (1999). Mixotrophy and nitrogen uptake by Pfiesteria piscicida (Dinophyceae). J. Phycol. 35,1430-1437. Libby, P., and Wheeler, P. (1997). Particulate and dissolved organic nitrogen in the central and eastern equatorial Pacific. Deep-Sea Res.44, 345-361. Lindell, M. J., Graneh, W, and Tranvik, L. J. (1995). Enhanced bacterial growth in response to photochemical transformation of dissolved organic matter. Limnol. Oceanogr 40,195-199. Lipschultz, F. (1995). Nitrogen-specific uptake rates of marine phytoplankton isolated from natural populations of particles by flow cytometry. Mar Ecol. Prog. Ser 123, 245-258, Lipschultz, F, and Owens, J. P. (1996). An assessment of nitrogen fixation as a source of nitrogen to the North Atlantic Ocean. Biogeochemistry 35, 261-274.

Dynamics of DON

241

Lisa, T., Piedras, P., Cardenas, J., and Pineda, M. (1995). Utilization of adenine and quanine as nitrogen sources by Chlamydomonas reinhardtii cells. Plant Cell Environ 18,583-588. Loh, A. N., and Bauer, J. E. (2000). Distribution, partitioning and fluxes of dissolved and particulate organic C, N and P in the eastern North Pacific and Southern Oceans. Deep-Sea Res. 147, 22872316. Lomas, M. W., and Gilbert, P. M. (1999). Temperature regulation of nitrate uptake: A novel hypothesis about nitrate uptake and reduction in cool-water diatoms. Limnol. Oceanogr. 44, 556-572. Lomas, M. W., Gilbert, P. M., Berg, G. M., and Burford, M. (1996). Characterization of nitrogen uptake by natural populations of Aureococcus anaphagejferens (Chrysophyceae) as a function of incubation duration, substrate concentration, light, and temperature. /. Phycol 32,907-916. Lomas, M. W., Trice, T. M., Gilbert, P. M., Bronk, D. A., and McCarthy, J. J. (in press). Temporal and spatial dynamics of urea concentrations in Chesapeake Bay: Biological versus physical forcing. Estuaries Lopez-Veneroni, D., and Cifuentes, L. A. (1992). Dissolved organic nitrogen distribution and transport in the continental shelf of the northwest Gulf of Mexico. "In Nutrient Enhanced Coastal Ocean Productivity," pp. 57-68. NOAA Lopez-Veneroni, D., and Cifuentes, L. A. (1994). Transport of dissolved organic nitrogen in Mississippi River Plume and Texas-Louisiana continental shelf near-shore waters. Estuaries 17, 796808. Lytle, C. R., and Perdue, E., M. (1981). Free, proteinaceous, and humic-bound amino acids in river water containing high concentrations of aquatic humics. Environ. Set Technol. 15,224-228. Malta, Y, and Yanada, M. (1990). Vertical distribution of total dissolved nitrogen and dissolved organic nitrogen in seawater. Geochem. J. 24, 245-254. Malone, T. C. (1980). Size-fractionated primary production. In "Primary Production in the Sea." Brookhaven Symp. Biol. 31, pp. 301-319. Plenum Press, New York. Mann, K. H. (1982). "Ecology of Coastal Waters. A System Approach," pp. 322. Blackwell Scientific, Oxford. Mantoura, R. F. C , Dickson, A., and Riley, J. P. (1978). The complexation of materials in natural waters. Estuarine Coastal Mar Sci. 6, 387-408. Marsot, P., Cembella, A. D., and Colombo, J. C. (1991). Intracellular and extracellular amino acid pools of the marine diatom Phaeodactylum tricomutum (Bacillariophyceae) grown on unemiched seawater in high cell density dialysis culture. /. Phycol 11,478-^91. Martinez, J., and Azam, F. (1993). Aminopeptidase activity in marine chroococcoid cyanobacteria. Appl. Environ Microbiol. 59, 3701-3707. Mason, C. A., and Engli, T. (1993). Dynamics of microbial growth in the decelerating and stationary phase of batch culture. In "Starvation in Bacteria" (S. Kjelleberg, Ed.), pp. 81-102. Plenum Press, New York. McCarthy, J. J. (1970). A urease method for urea in seawater. Limnol. Oceanogr 15, 309-313. McCarthy, M., Hedges, J., and Benner, R. (1996). Major biochemical composition of dissolved high molecular weight organic matter in seawater. Mar Chem. 55,281-297. McCarthy, M., Pratum, T, Hedges, J., and Benner, R. (1997). Chemical composition of dissolved organic nitrogen in the ocean. Nature 390,150-154. McCarthy, M. D., Hedges, J. L, and Benner, R. (1998). Major bacterial contribution to marine dissolved organic nitrogen. Science 281, 231-234. McKnight, D. M., and Aiken, G. R. (1998). Sources and age of aquatic humus. In "Aquatic Humic Substances" (D. O. Hessen and L. J. Tranvik, Eds.), pp. 9-39. Springer-Verlag, Berlin. Menzel, D. W., and Vaccaro, R. F. (1964). The measurement of dissolved organic and particulate carbon in seawater. Limnol. Oceanogr 9,138-142. Metzler, P., Gilbert, P., Gaeta, S., and Ludlam, J. (1997). New and regenerated production in the South Atlantic off Brazil. Deep-Sea Res. 44, 363-384.

242

Deborah A. Bronk

Metzler, P., Gilbert, P., Gaeta, S., and Ludlam, J. (2000). Contrasting effects of substrate and grazer manipulations on picoplankton In oceanic and coastal waters off Brazil. /. Plankton Res. 22, 77-90. Meybeck, M. (1993). C, N, P, and S in rivers: From sources to global inputs. In "Interactions of C, N, P, and S Biogeochemical Cycles and Global Change" (R. WoUast, F. T. Mackenzie, and L. Chou, Eds.), pp. 163-193. Springer-Verlag, Berlin. Middelboe, M., Borch, N. H., and Kirchman, D. L. (1995). Bacterial utilization of dissolved free amino acids, dissolved combined amino acids and ammonium in the Delaware Bay estuary: Effects of carbon and nitrogen limitation. Mar. Ecol Prog. Sen 128, 109-120. Middelboe, M., J0rgensen, N. O. G., and Kroer, N. (1996). Effects of viruses on nutrient turnover and growth efficiency of noninfected marine bacterioplankton. Appl. Environ Microbiol. 62,19911997. Middelburg, J. J., and Nieuwenhuize, J. (2000). Nitrogen uptake by heterotrophic bacteria and phytoplankton in the nitrate-rich Thames estuary. Mar Ecol. Prog. Sen 203,13-21. Miller, C. A., and Gilbert, P. M. (1998). Nitrogen excretion by the calanoid copepod Acartia tonsa: Results from mesocosm experiments. J. Plankton Res. 20, 1767-1780. Mopper, K., and Kieber, D. J. (2002). Photochemistry and cycUng of carbon, sulfur, nitrogen, and phosphorus. In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. Hansell and Carlson, C. A., Eds.), pp. 455-507. Academic Press, San Diego. Mopper, K., and Lindroth, P. (1982). Diel and depth variations in dissolved free amino acids and ammonium in the Baltic Sea determined by shipboard HPLC analysis. Limnol. Oceanogr. 27, 366-347. Mopper, K., Zhou, X., Kieber, R. J., Kieber, D. J., Sikorski, R. J., and Jones, R. D. (1991). Photochemical degradation of dissolved organic carbon and its impact on the oceanic carbon cycle. Nature 353, 60. Moran, M. A., and Hodson, R. E. (1994). Support of bacterioplankton production by dissolved humic substances from three marine environments. Man Ecol. Prog. Sen 110, 241-247. Moran, M. A., and Zepp, R. G. (2(XX)). UV radiation effects on microbes and microbial processes. In "Microbial Ecology of the Oceans" (D. L. Kirchman, Ed.), pp 201-228. Wiley-Liss, New York. Mortazavi, B., Iverson, R. L., Landing, W. M., Lewis, F. G., and Huang, W. (2000). Control of phytoplankton production and biomass in a river-dominated estuary: Apalachicola Bay, Florida, USA. Mar Ecol. Prog. Ser. 198, 19-31. MulhoUand, M. R., Gilbert, P M., Berg, G. M., Van Heukelem, L., Pantoja, S., and Lee, C. (1998). Extracellular amino acid oxidation by microplankton: A cross-system comparison. Aquat. Microb. Ecol. 15, 141-152. Miiller-Wegener, U. (1988). Interaction of humic substances with biota. In "Humic Substances and Their Role in the Environment" (F. H. Frimmel and R. F. Christman, Eds.), pp. 179-192. Wiley, New York. Miinster, U., and De Haan, H. (1998). The role of microbial extracellular enzymes in the transformation of dissolved organic matter in humic waters. In "Aquatic Humic Substances" (D. O. Hessen and L. J. Tranvik, Eds.), pp. 199-257. Springer-Verlag, Beriin. Murray, A. G. (1995). Phytoplankton exudation: Exploitation of the microbial loop as a defense against algal viruses. J. Plankton Res. 17, 1079-1094. Murray, A. G., and Eldridge, P. M. (1994). Marine viral ecology: Incorporation of bacteriophage into the microbial planktonic food web paradigm. J. Plankton Res. 16,627-641. Myklestad, S., Holm-Hansen, O., Varum, K. M., and Volcani, B. E. (1989). Rate of release of extracellular amino acids and carbohydrates from the marine diatom Chaetoceros affinis. J. Plankton Res. 11,763-773. Nagao, E, and Miyazaki, T. (1999). A modified ^^N tracer method and new calculation for estimating release of dissolved organic nitrogen by freshwater planktonic algae. Aquat. Microb. Ecol. 16, 309-314.

Dynamics of DON

243

Nagata, T. (2000). Production mechanisms of dissolved organic matter. In "Microbial Ecology of the Oceans" (D. L. Kirchman, Ed.), pp. 121-152. Wiley, New York. Nagata, T., and Kirchman, D. L. (1992). Release of macromolecular organic complexes by heterotrophic marine flagellates. Mar. Ecol Prog. Sen 83, 233-240. Nagata, T., and Kirchman, D. L. (1996). Bacterial degradation of protein adsorbed to model submicron particles in seawater. Man Ecol. Prog. Sen 132, 241-248. Nausch, M. (1996). Microbial activities on Trichodesmium colonies. Mar Ecol. Prog. Ser 141, 173181. Neilson, A. H., and Larsson, T. (1980). The utihzation of organic nitrogen for growth of algae: Physiological aspects. Physiol. Plant. 48, 542-553. Ogawa, H., Fukuda, R., and Koike, I. (1999). Vertical distributions of dissolved organic carbon and nitrogen in the Southern Ocean. Deep-Sea Res. 746, 1809-1826. Ohki, K. (1999). A possible role of temperate phage in the regulation of Trichodesmium biomass. In "Marine Cyanobacteria" (L. Charpy and A. W. D. Larkum, Eds.), Vol. 19, pp. 287-292. Bulletin de rinstitute Oceanographique, Monaco. Ohlendieck, U., Stuhr, A., and Siegmund, H. (2000). Nitrogen fixation by diazotrophic cyanobacteria in the Baltic Sea and transfer of the newly fixed nitrogen to picoplankton organisms. /. Mar Syst. 25, 213-219. O'Neil, J. M., Metzler, P. M., and Gilbert, P. M. (1996). Ingestion of ^^Ni -labelled Trichodesmium sp. and anmionium regeneration by the harpacticoid copepod Macrosetella gracilis. Mar Biol. 125, 89-96. O'Neil, J. M., and Roman, M. R. (1992). Grazers and associated organisms of Trichodesmium. In "Marine Pelagic Cyanobacteria: Trichodesmium and Other Diazotrophs" (E. J. Carpenter, D. G. Capone, and J. G. Renter, Eds.), pp. 61-73. Kluwer Academic, Dordrecht. Ouvemey, C. C , and Fuhrman, J. A. (2000). Marine planktonic archaea take up amino acids. Appl. Env. Microbiol. 66,4829^833 Paerl, H. W. (1988). Nuisance phytoplankton blooms in coastal, estuarine, and inland waters. Limnol. Oceanogr 33, 823-847. Page, A. L., Miller, R. H., and Keeney, D. R. (1982). "Methods of Soil Analysis" pp. 1159. Soil Science Society of America, Inc., Madison. Palenik, B., and Henson, S. E. (1997). The use of amides and other organic nitrogen sources by the phytoplankton Emiliana huxleyi. Limnol. Oceanogr 42, 1544-1551. Palenik, B., and Koke, J. A. (1995). Characterization of a nitrogen-regulated protein identified by cell surface biotinylation of a marine phytoplankton. Appl. Env. Microbiol. 61, 3311-3315. Palenik, B., and Morel, F. M. M. (1990a). Amino acid utilization by marine phytoplankton: A novel mechanism. Limnol. Oceanogr 35,260-269. Palenik, B., and Morel, F. M. M. (1990b). Comparison of cell-surface L-amino acid oxidases form several marine phytoplankton. Mar Ecol. Prog. Ser 59,195-201. Palenik, B., and Morel, F. M. M. (1991). Amine oxidases of marine phytoplankton. App. Environ. Microbiol. 57, 2440-2443. Pantoja, S., and Lee, C. (1994). Cell-surface oxidation of amino acids in seawater. Limnol. Oceanogr 39,1718-1726. Pantoja, S., and Lee, C. (1999). Peptide decomposition by extracellular hydrolysis in coastal seawater and salt marsh sediment. Mar Chem. 63,273-291. Pantoja, S., Lee, C , and Marecek, J. F. (1997). Hydrolysis of peptides in seawater and sediment. Mar Chem. 5 7 , 2 5 ^ 0 . Parsons, T. R., Malta, Y., and Lalli, C. (1984). "A Manual of Chemical and Biological Methods for Seawater Analysis." Pergamon Press, Oxford. Paul, J. H., DeFlaun, M. R, Jeffrey, W. H., and David, A. W. (1988). Seasonal and diel variability in dissolved DNA and in microbial biomass and activity in a subtropical estuary. Appl. Environ. Microbiol. 54, llS-121.

244

Deborah A. Bronk

Paul, J. H., Jeffrey, W. H., and DeFlaun, M. F. (1987). Dynamics of extracellular DNA in the marine environment. Appl. Environ. Microbiol. 53,170-179. Poulet, S. A., and Martin-Jezequel, V. (1983). Relationships between dissolved free amino acids, chemical composition and growth of the marine diatom Chaetoceros debile. Mar. Biol. 11,93-100. Poutanen, E. L., and Morris, R. J. (1985). Comparison of the structures of humic acids from marine sediments and degraded field diatoms by ^^C— and ^H-NMR spectroscopy. Mar Chem. 17,115126. Prakash, A. (1971). Terrigenous organic matter and coastal phytoplankton fertihty. In "FertiHty of the Sea" (J. D. Costlow, Ed.), pp. 351-368. Gordon and Breach, New York. Prakash, A., M.A., R,, Jensen, A., and Rao Subba, D. V. (1973). Influence of humic substances on the growth of marine phytoplankton: Diatoms. Limnol. Oceanogr 18,516-524. Price, N., and Harrison, P. (1987). Comparison of methods for the analysis of dissolved urea in seawater. Man Biol. 94, 301-311. Price, N. M., and Harrison, P. J. (1988). Urea uptake by Sargasso Sea phytoplankton: Saturated and in situ uptake rates. Deep-Sea Res. 35,1579-1593. Probyn, T. A. (1985). Nitrogen uptake by size-fractionated phytoplankton populations in the southern Benguela upweUing systems. Mar Ecol. Prog. Sen 22, 249-258. Probyn, T. A., Waldron, H. N., and James, A. G. (1990). Size-fractionated measurements of nitrogen uptake in ages upwelled waters: Implications for pelagic food webs. Limnol. Oceanogr 35, 202210. Pujo-Pay, M., Conan, P., and Raimbault, P. (1997). Excretion of dissolved organic nitrogen my phytoplankton assessed by wet oxidation and N-15 tracer procedures. Mar Ecol. Prog. Ser 153,99-111. Raimbault, P., Slawyk, G., Boudjellal, B., Coatanoan, C , Conan, P., Coste, B., Carcia, N., Moutin, T., and Pujo-Pay, M. (1999). Carbon and nitrogen uptake and export in the equatorial Pacific at 150°W: Evidence for an efficient regenerated production cycle. J. Geophys. Res. 104, 3341-3356. Rice, D. L. (1982). The detritus nitrogen problem: New observations and perspectives from organic geochemistry. Mar Ecol. Prog. Ser 9, 153-162. Rich, J., Gosselin, M., Sherr, E., Sherr, B., and Kirchman, D. L. (1997). High bacterial production, uptake and concentrations of dissolved organic matter in the Central Arctic ocean. Deep-Sea Res. 7/44,1645-1663. Riegman, R., Flameling, I. A., and Noordeloos, A. A. M. (1998). Size-fractionated uptake of ammonium, nitrate and urea and phytoplankton growth in the North Sea during spring 1994. Mar Ecol. Prog. Ser 173, 85-94. Riegman, R., and Noordeloos, A. A. M. (1998). Size-fractionated uptake of nitrogenous nutrients and carbon by phytoplankton in the North Sea during sunmier 1994. Mar Ecol. Prog. Ser 173,95-106. Roy, S., Harris, R. P., and Poulet, S. A. (1989). Inefficient feeding by Calanus helgolandicus and Temora longicomis on Coscinodiscus wailesii : Quantitative estimation using chlorophyll-type pigments and effects on dissolved free amino acids. Mar Ecol. Prog. Ser 52, 145-153. Sakano, S., and Kamatani, A. (1992). Determination of dissolved nucleic acids in seawater by the fluorescence dye, ethidium bromide. Mar Chem. 37, 239-255. Sakshaug, E., and Holm-Hansen, O. (1977). Chemical composition of Skeletonema costatus (Grev.) Cleve and Pavlova (Monochrysis) lutheri (Droop) Green as a function of nitrate-, phosphate-, and iron-limited growth. /. Exp. Mar Biol. Ecol. 29, 1-34. Sanders, R., and Jickells, T (2000). Total organic nutrients in Drake Passage. Deep-Sea Res. I 47, 997-1014. Schnef, E., and Elbracter, M. (1992). Nutritional strategies of dinoglagellates. A review with emphasis on cell biological aspects. J. Protistol. 28, 3-24 Schnitzer, M. (1985). Nature of nitrogen in humic substances. In "Humic Substances in Soil, Sediment, and Water" (G. R. Aiken, D. M. McKnight, and R. L. Wershaw, Eds.), pp. 303-328. Wiley, New York.

Dynamics of DON

245

Schuster, S., Arrieta, J. M., and Hemdl, G. J. (1998). Adsorption of dissolved free amino acids on colloidal DOM enhances colloidal DOM utilization but reduces amino acid uptake by orders of magnitude in marine bacterioplankton. Mar. EcoL Prog. Sen 166,99-108. Seitzinger, S., and Sanders, R. (1997). Contribution of dissolved organic nitrogen from rivers to estuarine eutrophication. Man Ecol. Prog. Sen 159,1-12. Seitzinger, S., and Sanders, R. (1999). Atmospheric inputs of dissolved organic nitrogen stimulate estuarine bacteria and phytoplankton. Limnol Oceanogr. 44,721-730. Sellner, K. G. (1992). Trophodynamics of marine cyanobacteria blooms. In "Marine Pelagic Cyanobacteria: Trichodesmium and Other Diazotrophs" (E. J. Carpenter, D. G. Capone, and J. G. Renter, Eds.), pp. 75-95. Kluwer Academic Publishers, The Netherlands. Sellner, K. G., and Nealley, E. W. (1997). Diel fluctuations in dissolved free amino acids and monosaccharides in Chesapeake Bay dinoflagellate blooms. Mar Chem. 56,193-200. Sharp, J. H. (1973). Total organic carbon in seawater—comparison of measurements using persulfate oxidation and high temperature combustion. Man Chem. 1, 211-229. Sharp, J. H. (1983). The distribution of inorganic nitrogen and dissolved and particulate organic nitrogen in the sea. In "Nitrogen in the Marine Environment" (D. G. Capone and E. J. Carpenter, Eds.), pp. 101-120. Plenum Press, New York. Sharp, J. H. (2002). Analytical Methods for Total DOM Pool In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. Hansell and Carlson, C. A., Eds.), pp. 35-58. Academic Press, San Diego. Sharp, J. H., Rinker, K. R., Savidge, K. B., Abell, J., Benaim, J. Y., Bronk, D. A., Burdige, D. J., Cauwet, G., Chen, W., Doval, M. D., Hansell, D., Hopkinson, C , Kattner, G., Kaumeyer, N., McGlathery, K. J., Merriam, J., Morley, J., Nagel, K., Ogawa, H., Pollard, C , Raimbault, P., Seitzinger, S., Spyres, G., Tirendi, K, Walsh, T. W. and Wong, C. S. (in press). A preliminary methods comparison for measurement of dissolved organic nitrogen in seawater. Man Chem. Shibata, A., Kogure, K., Koike, L, and Ohwada, K. (1997). Formation of submicron colloidal particles from marine bacteria by viral infection. Man Ecol Prog. Sen 155, 303-307. Shimizu, Y, Watanabe, N., and Wrensford, G. (1995). Biosynthesis of brevetoxins and heterotrophic metabolism in Gymnodinium breve. In "Harmful Marine Algal Blooms" (P. Lassus, G. Arzul, E. Erard, P. Gentien, P., and C. Marcaillou, Eds.), Lavoisier, Intercept, Ltd. Sigman, S. D., Altabet, M. A., Michner, R., McCorkle, D. C , Fry, B., and Holmes, R. M. (1997). Natural abundance-level measurement of the nitrogen isotopic composition of oceanic nitrate: An adaptation of the ammonia diffusion method. Man Chem. SI, 227-242. Slawyk, G., and Raimbault, P. (1995). Simple procedure for simultaneous recovery of dissolved inorganic and organic nitrogen in ^^N-tracer experiments and improving the isotopic mass balance. Man Ecol. Prog. Sen V2A, 289-299. Slawyk, G., Raimbault, P., and Garcia, N. (1998). Measuring gross uptake of ^^N-labeled nitrogen by marine phytoplankton without particulate matter collection: Evidence for low ^^N losses to the dissolved organic nitrogen pool. Limnol. Oceanogn 43,1734-1739. Slawyk, G., Raimbault, P., and Garcia, N. (2000). Use of ^^N to measure dissolved organic nitrogen release by marine phytoplankton. Limnol. Oceanogn 45,1884-1886. Smith, D. C , Simon, M., AUdredge, A. L., and Azam, F. (1992). Intense hydrolytic enzyme activity on marine aggregates and implications for rapid particle dissolution. Nature 359,139-141. Smith, S. v., HoUibaugh, J. T., DoUan, S. J., and Vink, S. (1991). Tomales Bay metabohsm: C-N-P stoichiometry and ecosystem heterotrophy at the land-sea interface. Estuarine Coastal Shelf Sci. 33, 223-257. S0ndergaard, M., WiUiams, P. J. le. B., Cauwet, G., Riemann, B., Robinson, C , Terzic, S., Woodward, E. M. S., and Worm, J. (2000). Net accumulation and flux of dissolved organic carbon and dissolved organic nitrogen in marine plankton communities. Limnol. Oceanogn 45,1097-1 111. Steidinger, K. A., Vargo, G. A., Tester, P. A., and Tomas, C. R. (1998). Bloom dynamics and physiology of Gymnodinium breve with emphasis on the Gulf of Mexico. In "Physiological Ecology of Harmful

246

Deborah A. Bronk

Algal Blooms" (D. M. Anderson, A. D. Cembella, and G. M. Hallegraeff, Eds.), Springer-Verlag, Beriin Stepanauskas, R., Edling, H., and Tranvik, L. J. (1999b). Differential dissolved organic nitrogen availability and bacterial aminopeptidase activity in limnic and marine waters. Microb. Ecol. 38, 264272. Stepanauskas, R., j0rgensen, N. O. G., Eigaard, O. R., and Zvikas, A. (in press). Riverine nutrients in the Baltic Sea. Ecological Monographs. Stepanauskas, R., Laudon, H., and J0rgensen, N. O. G. (2000). High DON bioavailability in boreal streams during a spring flood. Limnol Oceanogr. 45, 1298-1307. Stepanauskas, R., Leonardson, L., and Tranvik, L. J. (1999a). Bioavailability of wetland-derived DON to freshwater and marine bacterioplankton. Limnol Oceanogr. 44,1477-1485. Stevenson, F. J. (1994). "Humus Chemistry." Wiley, New York. Strom, S. L. (2000). Bacterivory: Interactions between bacteria and their grazers. In "Microbial Ecology of the Oceans" (Kirchman, D. L., Ed.), pp. 351-386. Wiley-Liss, New York. Strom, S. L., Benner, R., Ziegler, S., and Dagg, M. J. (1997). Planktonic grazers are a potentially important source of marine dissolved organic carbon. Limnol. Oceanogr 42, 1364-1374. Stuermer, D. H., Peters, K. E., and Kaplan, I. R. (1978). Source indicators of humic substances and proto-kerogen: Stable isotope ratios, elemental compositions, and electron spin resonance spectra. Geochim. Cosmochim. Acta 42, 989-997. Suttle, C. A. (1994). The significance of viruses to mortality in aquatic microbial communities. Microb. Ecol. 28, 237-243. Suzuki, Y, and Sugimura, Y (1985). A catalytic oxidation method for the determination of total nitrogen dissolved in seawater. Mar Chem. 16, 83-97. Tanmiinen, T, and Irmisch, A. (1996). Urea uptake kinetics of a midsummer planktonic community on the SW coast of Finland. Mar Ecol. Prog. Ser 130, 201-211. Tamminen, T, and Seppala, J. (1999). Nutrient pools, transformations, ratios, and hmitations in the Gulf of Riga, the Baltic Sea, during four successional stages. J. Mar Systems 23, 83-106. Tanoue, E., Ishii, M., and Midorikawa, T. (1996). Discrete dissolved and particulate proteins in oceanic waters. Limnol. Oceanogr 41,1334-1343. Tanoue, E., Nishiyama, S., Kamo, M., and Tsugita, A. (1995). Bacterial membranes: Possible source of a major dissolved protein in seawater. Geochim. Cosmochim. Acta 59, 2643-2648. Therkildsen, M., Isaksen, M. F , and Lonstein, B. A. (1997). Urea production by the marine bacteria Delaya venusta and Pseudomonas stutzeri grown in a minimal medium. Aquat. Microb. Ecol. 13, 213-217. Thurman, E. M. (1985). "Organic Geochemistry of Natural Waters." Niyhoff/Junk, Boston. Tobias, C. R., Harvey, J. W., and Anderson, I. C. (2001). Quantifying groundwater discharge through fringing wetlands to estuaries: Seasonal variability, methods comparison, and implications for wetland-estuary exchange. Limnol. Oceanogr 46, 604-615. Toggweiler, J. R. (1989). Is the downward dissolved organic matter (DOM) flux important in carbon transport? In "Productivity in the Ocean: Past and Present" (W H. Berger, V. S. Smetacek, and G. Wefer, Eds.), pp. 65-83. Wiley, New York. Toledo, A. P. P., D'Aquino, V. A., and Tundisi, J. G. (1982). Influence of humic acid on growth and tolerance to cupric ions in Melosira italica (subsp. antarctica). Hydrobiology 87, 247-254. Tsugita, A., Uchida, T, Mewes, H. W, and Atake, T. (1987). A rapid vapor-phase acid (hydrochloric acid and trifluoroacetic acid) hydrolysis for peptide and protein. J. Biochem. 102,1593-1597. Tupas, L., and Koike, I. (1990). Amino acid and ammonium utiUzation by heterotrophic marine bacteria grown in enriched seawater. Limnol. Oceanogr 35,1145-1155. Turk, v., Rehnstam, A.-S., Lundberg, E., and Hagstrom, A. (1992). Release of bacterial DNA by marine nanoflagellates, an intermediate step in phosphorus regeneration. Appl. Env. Microb. 58, 3744-3750.

Dynamics of DON

247

Urban-Rich, J. (1999). Release of dissolved organic carbon from copepod fecal pellets in the Greenland Sea. /. Exp. Mar. Biol. Ecol. 232,107-124. Valderrama, J. C. (1981). The simultaneous analysis of total nitrogen and total phosphorus in natural waters. Mar. Chem. 10,109-122. Valiela, I., Costa, J., Foreman, K., Teal, J. M., Howes, B., and Aubrey, D. (1990), Transport of groundwater-bome nutrients from watersheds and their effects on coastal waters. Biogeochemistry 10, 177-197. Valiela, I., and Teal, J. M. (1979). The nitrogen budget of a salt marsh ecosystem. Nature 280,652-656. Vidal, M., Duarte, C. M., and Agusti, S. (1999). Dissolved organic nitrogen and phosphorus pools and fluxes in the central Atlantic Ocean. Limnol. Oceanogn 44,106-115. Villareal, T. A., Altabet, M. A., and Culver-Rymsza, K. (1993). Nitrogen transport by vertically migrating diatom mats in the North Pacific Ocean. Nature 363,709-712. Waldron, H. N., Attwood, C. G., Probyn, T. A., and Lucas, M. I. (1995). Nitrogen dynamics in the Bellingshausen Sea during the Austral spring of 1992. Deep-Sea Res. II42, 1253-1276. Wang, W., Tarr, M. A., Bianchi, T. S., and Engelhaupt, E. (2000). Ammonium photoproduction from aquatic humic and colloidal matter. Aquat. Geochem. 6,275-292. Ward, B. B., and Bronk, D. A. (2001). Net nitrogen uptake and DON release in surface waters: Importance of trophic interactions implied from size fractionation experiments. Mar. Ecol. Prog. Ser. 2\%n-2A. Weinbauer, M. G., and Peduzzi, P. (1995). Effect of virus-rich high molecular weight concentrates of seawater on the dynamics of dissolved amino acids and carbohydrates. Mar Ecol. Prog. Ser 127, 245-253. Wheeler, P. A., and Hellebust, J. A. (1981). Uptake and concentration of alkylamines by a marine diatom. Plant. Physiol. 67, 367-372. Wheeler, P. A., and Kirchman, D. L. (1986). Utilization of inorganic and organic nitrogen by bacteria in marine systems. Limnol. Oceanogr 31,998-1009. Wheeler, P. A., Watkins, J. M., and Hansing, R. L. (1997). Nutrients, organic carbon and organic nitrogen in the upper water column of the Arctic Ocean: Implications for the sources of dissolved organic carbon. Deep-Sea Res. 7/44,1571-1592. Williams, P. J. le. B. (2000). Heterotrophic bacteria and the dynamics of dissolved organic material. In "Microbial Ecology of the Oceans" (Kirchman, D. L., Ed), pp. 153-200. Wiley-Liss, New York. Williams, R., and Poulet, S. A. (1986). Relationships between the zooplankton, phytoplankton, particulate matter, and dissolved free amino acids in the Celtic Sea. Man Biol. 90, 279-284. Wommack, K. E., and Colwell, R. R. (2000). Virioplankton: Viruses in aquatic ecosystems. Microbiol. Mol. Biol. Rev. 64,69-114.

This Page Intentionally Left Blank

Chapter 6

Dynamics of DOP D. M. Karl and K. M. Bjorkman Department of Oceanography School of Ocean and Earth Science and Technology University of Hawaii Honolulu, Hawaii I. Introduction II. Terms, Definitions, and Concentration Units III. The Early Years of Pelagic Marine P-cycle Research (1884-1955) IV. The Pelagic Marine P-cycle: Key Pools and Processes V. Sampling, Incubation, Storage, and Analytical Considerations A. Sampling B. Use of Isotopic Tracers in P-cycle Research C. Sample Processing, Preservation, and Storage D. Detection of P/ and P-containing Compounds in Seawater E. Analytical Interferences in SRP and TDP Estimation VI. DOP in the Sea: Variations in Space A. Regional and Depth Variations in DOP B. DOP Concentrations in the Deep Sea

Biogeochemistry of Marine Dissolved Organic Matter Copyright 2002, Elsevier Science (USA). All rights reserved.

C. C:N:P Stoichiometry of Dissolved and Particulate Matter Pools VII. DOP in the Sea: Variations in Time A. English Channel B. North Pacific Subtropical Gyre VIII. DOP Pool Characterization A. Molecular Weight Characterization of the DOP Pool B. DOP Pool Characterization by Enzymatic Characterization C. DOP Pool Characterization by3ipNMR D. DOP Pool Characterization by Partial Photochemical Oxidation E. Direct Measurement of DOP Compounds F. Biologically Available P G. DOP: The "Majority" View IX. DOP Production, Utilization, and Remineralization A. DOP Production and Remineralization

249

250

Karl and Bjorkman B. Direct Utilization of DOP C. Enzymes as P-cycle Facilitators D. DOP Interactions with Light and Suspended Minerals

X. Conclusions and Prospectus References

I. INTRODUCTION Phosphorus (P) is an essential macronutrient for all living organisms; life is truly built around P (deDuve, 1991). In the sea, P exists in both dissolved and particulate pools with inorganic and organic forms. The uptake, remineralization and physical and biological exchanges among these various pools are the essential components of the marine P cycle (Fig. 1). Compared to the much more comprehensive investigations of carbon (C) and nitrogen (N) dynamics in the sea, P pool inventories and fluxes are less well documented though no less important. During cell growth, P is incorporated into a broad spectrum of organic compounds with vital functions including structure, metabolism, and regulation. In time, selected P-containing organic compounds are lost from the cells to the surrounding environment by combined exudation and excretion processes. When cells turn over, whether by death/autolysis, grazing, parasitism, or viral infection, there is an enhanced release of intracellular P-containing compounds as both dissolved and particulate organic matter (DOM and POM, respectively). In this broad view, dissolved organic P (DOP) is simply the intermediate between inorganic P (P/) uptake and Fi regeneration (Fig. 1). For this and other ecological and analytical interdependencies of P/ and DOP, it is impossible to isolate DOP from the remainder of the marine P cycle. It is also imperative to emphasize that the production and cycling of P-containing compounds are inextricably linked to C and N dynamics by virtue of the fact that marine DOM and POM pools include many compounds that contain both C and P (e.g., phospholipids, sugar phosphates and selected vitamins and phosphonates) or C, N, and P (e.g., nucleotides, nucleic acids, and selected phosphonates; see Figs. 2 and 3 and Table I). It is, therefore, inappropriate to consider DOP as separate from dissolved organic C (DOC) and dissolved organic N (DON) or to view the P-cycle in any similar biogeochemical isolation. This review will take a holistic approach to the marine P-cycle with an emphasis on the production and turnover of P-containing and N-and-P-containing dissolved organic matter (i.e., DOC-P and DOC-N-P, hereafter collectively referred to as DOP). By design, this chapter will focus on the pelagic environment, especially the open sea. Investigation of the marine sedimentary P cycle is further complicated by the presence of numerous poorly defined P reservoirs (e.g., Ruttenberg, 1992;

Dynamics ofDOP

251

-ft'-

Atm

/ \ ^ y'v. / I \

Deposition (wet and dry)

circulation processes DOP/

Extreme photolysis

Pi

PH,

DOP

VO,

+

P/

inorganic poly-P/ / tiydroiysis

diffusion, mixing, active transport

Ecto

4r

diffusion, mixing, active transport

Particulate P

\

Low Density Upward Flux

Export Flux (Gravitational and active migrations)

Figure 1 The open-ocean P cycle showing the various sources and sinks of inorganic and organic P, including biotic and abiotic interconversions. The large rectangle in the center represents the upper water column TDP pool composed of Ft, inorganic polyphosphate, and a broad spectrum of largely uncharacterized DOP compounds. Ectoenzymatic activity (Ecto) is critical for microbial assimilation of selected TDP compounds. Particulate P, which includes all viable microorganisms, sustains the P cycle by assimilating and regenerating Pi, producing and hydrolyzing selected non-P/ P, especially DOP compounds, and supporting net particulate matter production and export. Atmospheric deposition, horizontal transport, and the upward flux of low-density organic P compounds are generally poorly constrained processes in most marine habitats. Phosphine (PH3), shown at the right, is the most reduced form of P in the biosphere and is generally negligible except under very unusual, highly reduced conditions. Redrawn from Karl and Bjorkman (2001).

Anderson and Delaney, 2000) which precludes a straightforward determination of P inventories and fluxes. While the majority of P-cycle processes occur throughout the world's oceans, net DOM/POM production is enhanced in the euphotic zone (e.g., the upper 0-100 m of the water column) while net remineralization of DOM/POM generally occurs at greater water depths. This vertical stratification of the marine P cycle (as well as C and N cycles) is an important factor which ultimately controls the distributions and abundances of microbial biomass and rates of global ocean biomass production, and greatly impacts the sources, sinks, and, most likely, chemical composition of marine DOM.

Karl and Bprkman

252 CARBON I CO, "^°^ COi hydrocarbons monosaccharides fatty acids vitD amino acids '

/ PPi " /1 PPPil

hexose-P triose-P RuBP PEP phospholipids

\

PH3\

\

^ selected

nucleotides nucleic acids

amino sugars

teichoic acids

chlorophyll a

selected phosphonates

protein

vit Bi and B12 humic/fuMc adds

\

urea

chitin peptidoglycan

\ NO-

|No; INH; /NgO

lipids polysaccharides cellulose/starch

Figure 2 Bar and shield representation of dissolved matter in seawater showing the intersection of C, N, and P compound classes. For example, dissolved P can exist in a variety of inorganic P forms (outside portion of the shield) or as DOC-P and DOC-N-P compounds. Likewise, C and N have both unique and intersecting pools. Compound symbols include: Fi, orthophosphate; PFi, pyrophosphate (pyro-PO; PPP/, inorganic polyphosphate (poly-P/); PH3, phosphine; RuBP, ribulose bisphosphate; PEP, phosphoenolpyruvate; N2, nitrogen; N02~, nitrite; NOg", nitrate; NH4'^, ammonium; and N2O, nitrous oxide.

We will present selected information on DOP formation, distribution and turnover in the sea building upon several previous, authoritative reviews by Duursma (1960), Armstrong (1965), Comer and Davies (1971), and Benitez-Nelson (2000) on various aspects of the marine P cycle, as well as nearly one century of field and laboratory research on this subject. For reasons already mentioned, it is impossible to discuss DOP in any useful ecological framework without also considering other DOM/POM pools and related biogeochemical processes. Although dissolved inorganic P concentrations (typically reported as soluble reactive phosphorus or SRP) are routinely measured in physical, chemical, and biological studies of the marine environment, estimates of total P (i.e., the sum of reactive and nonreactive forms of dissolved P, also called total dissolved P or TDP) are rare, despite the existence, for over 50 years, of reliable analytical methods. Although TDP was included as a core measurement during the International Geophysical Year (IGY) Atlantic Basin hydrographic survey of 1957-1958 (McGill, 1963), none

Dynamics ofDOP

253

of the "modem" oceanographic sampling programs, including Geochemical Ocean Sections (GEOSECS) and World Ocean Circulation Experiment (WOCE) included TDP as a core measurement. Even the Joint Global Ocean Flux Study (JGOFS) program, which sponsored regional-scale field studies of ocean biogeochemistry, mostly ignored P-cycle processes. Consequently the extant database of high-quality, paired SRP and TDP in the world's oceans is relatively sparse in comparison to the global coverage of SRP.

II. TERMS, DEFINITIONS, AND CONCENTRATION UNITS The total phosphorus (TP) fraction in seawater is divided, unequally, among particulate P (PP) and TDP fractions (TP = PP + TDP); both fractions contain inorganic and organic P derivatives. In most open ocean marine environments, the TDP pool greatly exceeds the PP pool, but it is the biogenic PP pool (i.e., cells or living biomass) that ultimately produces and remineralizes DOP, thereby sustaining the marine P cycle. The inorganic forms of P consist mostly of orthophosphoric acid (in seawater at a salinity of 33%^, 20°C, and pH 8.0 as 1% H2PO4- / 87% HPO/~/12% P04^-; Kester and Pytkowicz, 1967), pyrophosphate (P2O/"; hereafter abbreviated pyro-PO, and other condensed cyclic (metaphosphate) and linear (polyphosphate) polymers of various molecular weights (hereafter abbreviated poly-P/). The condensed phosphates can exist in the dissolved, colloidal and particulate matter fractions of seawater, whereas P/ and pyro-P/ are mostly contained in the truly dissolved fraction or within intracellular pools. Of these various inorganic forms, only P/ is quantitatively detected by the standard molybdenum blue assay procedure (see Section V.D for more information on reaction specificity). Therefore the measurements of pyro-P/ and poly-P/ pools require sample hydrolysis to yield reactive Fi, The organic-P fractions include primarily monomeric and polymeric phosphate esters (C-O-P bonded compounds), phosphonates (C-P bonded compounds), and organic condensed phosphates (Table I and Fig. 3). Among the ester-linked DOP compounds, both phosphomonoesters and phosphodiesters are present (Table I); each compound has unique chemical and physical properties, and each has characteristic phosphohydrolytic enzyme susceptibility. Numerous compound classes (e.g., nucleotides, nucleic acids, phospholipids, phosphoproteins, sugar phosphates, phosphoamides, vitamins) have been detected in seawater and these will be discussed in subsequent sections. Oxidative destruction of the associated organic matter is generally required to convert organic-P to reactive Fi, although certain compound classes are partially hydrolyzed during Fi analysis and thus may contribute to an overestimation of the true Fi concentration. For this reason, the standard molybdenum blue assay measures an operationally defined pool.

BOND TYPE C-O-P (Monoester)

" \ /

Example: Glucose-6-phosphate

HO

C-0-P-O-C (Diester)

NH,

o=p—o-

Example: Ribonucleic acid

H I

f H

O

C-P (Phosphonate)

OH

OH

Example: Phosphonoformic acid OH

C-0-P-O-P-O-P (Polyphosphate monoester) Example: Adenosine-5'-triphosphate

HO

P*^^^^

P*^^-^

OH

OH

OH H I

( H

Dynamics ofDOP

255

soluble reactive P (SRP), and the difference between TDP (i.e., equal to SRP following sample hydrolysis) and the initial SRP value has been termed the soluble nonreactive P (SNP) pool. Although SRP is often equated to P/, in reaUty SRP only sets an upper constraint on P/. Depending upon oxidation/hydrolysis conditions that are used for analysis, the SNP pool includes organic-P, pyro-P/, and poly-P/. Consequentiy, SNP concentration is technically not equal to DOP due to the two independent conditions: SRP>P/ and SNP>DOP. This may have important analytical and ecological implications as discussed in subsequent sections. P in seawater can also be characterized by origin (e.g., biogenic or lithogenic) or by physical characteristics (e.g., molecular weight or photolytic lability). Because many different forms can be used as P sources for marine microorganisms, albeit at variable rates and efficiencies, the most ecologically relevant fraction is biologically available P (BAP) pool. Ideally, BAP consisting of both Ft plus the biolabile fraction of the SNP pool should be measured to constrain oceanic P cyclefluxes,but routine analytical methods do not exist. While it might be argued, a priori, that SRP measurements by the Murphy and Riley (1962) procedure place a lower bound on BAP, because both Fi and acid-labile DOP must be biologically available, this may not always be the case. Fi contained in colloidal associations or adsorbed to nanoparticles would assay as part of the SRP pool but might be unavailable for uptake under ambient conditions. In all likelihood only microbioassay analysis can provide an accurate estimate of BAP (see Section VIII.F). Suffice it to say, we are still lacking a comprehensive chemical description of dissolved P in seawater (see Section VIII). The measurement of TDP is also operationally defined; typically a highintensity ultraviolet (UV) photooxidation (Armstrong et al, 1966) or hightemperature wet chemical oxidation (Menzel and Corwin, 1965) or a combined (Ridal and Moore, 1990) pretreatment is used to convert SNP to Fi for subsequent analysis by the standard molybdenum blue assay. However, it is well known that certain P-containing compounds (e.g., poly-P/, nucleotide di- and triphosphates) are not quantitatively recovered by standard UV photooxidation procedures; neither method quantitatively recovers P from all phosphonate compounds. Depending upon the methods used, the difference between the measurement of TDP and either Fi or SRP can be termed SNP (i.e., SNP = [TDP]-[SRP]) or non-Pf P (non-P/ P = [TDP]-[P/]), where SNP ^ N-Fi F (Thomson-Bulldis and Karl, 1998). As emphasized previously, there is no a priori relationship between these operationally defined pools and the more ecologically relevant BAP pool. Although several SNP compound classes have been reported to exist in seawater, including

Figure 3 Selected structures of representative DOP pool compound classes with specific examples. Not shown here are the less well known classes such as phosphoramidates (N-P-bonded) or phosphorothionates (S-P-bonded) compounds that could also be present in cells and in seawater.

Table I Selected Inorganic and Organic P Compounds Either Known to Be or Likely to Be Present in Seawater Compound Monophosphate esters Ribose-5-phosphoric acid (R-5-P) Phospho(enol)pyruvic acid (PEP) Glyceraldehyde 3-phosphoric acid (G-3-P) Glycerophosphoric acid (Gly-3-P) Creatine phosphoric acid (CP) Glucose-6-phosphoric acid (Glu-6-P) Ribulose-l,5-bisphosphoric acid (RuBP) Fructose-1,6-diphosphoric acid (F-1,6-DP) Phosphoserine (PS) Nucleotides and derivatives Adenosine 5'-triphosphoric acid (ATP) Uridylic acid (UMP) Uridine diphosphate glucose (UDPG) Guanosine 5'-diphosphate 3'-diphosphate or "magic spot" (ppGpp) Pyridoxal 5-monophosphoric acid (PyMP) Nicotinamide adenine dinucleotide phosphate (NADP) Ribonucleic acid (RNA) Deoxyribonucleic acid (DNA) Inositohexaphosphoric acid or phytic acid (PA) Vitamins Thiamine pyrophosphate (vitamin Bi) Riboflavine 5'-phosphate (vitamin B2-P) Cyanocobalamin (vitamin B12)

Chemical formula (molecular weight)

P (% by weight)

Molar C:N:P

CsHiiOgP (230.12) C3H5O6P (168) C3H7O6P (170.1) C3H9O6P (172.1) C4H10N3O5P (211.1) C6H13O9P (260.14) C5H60nP2 (304) C6H14O12P2 (340.1) C3H8NO6P (185.1)

13.5

5:—:1

18.5

3:—:1

18.2

3:—:1

18.0

3:—:1

14.7

4:3:1

11.9

6:—:1

20.4

2.5:—:1

18.2

3:—:1

16.7

3:1:1

18.3

3.3:1.7:1

9.6

9:2:1

10.9

7.5:1:1

20.6

2.5:1.25:1

12.5

8:1:1

9.4

11:3:1

-9.2% -9.5% 28.2

-9.5:4:1 -10:4:1 1:—:1

14.6

6:2:1

6.8

17:4:1

2.3

63:14:1

C10H16N5O13P3 (507.2) C9H13N2O9P (324.19) C15H24N2O17P2 (566.3) C10H17N5O17P4 (603) CgHioNOeP (247.2) C22H28N2O14N6P2 (662) Variable Variable C6H18O24P6 (660.1) C12H19N4O7P2S (425) C17H21N4O9P (456.3) C63H88C0N14O14P (1355.42)

(Continues)

257

Dynamics ofDOP Table I Compound Phosphonates Methylphosphonic acid (MPn) Phosphonoformic acid (FPn) 2-aminoethylphosphonic acid (2-AEPn) Other compounds/compound classes Marine fulvic acid'^ (MFA) Marine humic acid'' (MHA) Phospholipids (PL) Malathion (Mai) "Redfield" plankton

(Continued)

Chemical formula (molecular weight)

P (% by weight)

Molar C:N:P

CH5O3P (96) CO5PH3 (126) C2H8NO4P (141)

32.3

1:—:1

24.6

1:—:1

22.0

2:1:1

0.4-0.8 0.1-0.2 300:—:1 -40:1:1 9:—:1

1-3

106:16:1

Variable Variable Variable C9H16O5PS (267) Variable

^Marine FA and HA are operationally defined fractions, thus their composition may vary with source. These values are from Nissenbaum (1979).

poly-P/ (Solorzano and Strickland, 1968), nucleotides (Azam and Hodson, 1977; Nawrocki and Karl, 1989), nucleic acids (DeFlaun et al, 1986; Karl and Bailiff, 1989), and monophosphate esters (Strickland and Solorzano, 1966), the SNP pool in seawater remains largely uncharacterized. The earliest reports of P/ and TDP in seawater, prior to approximately 1930, all reported P as milligrams of phosphorus pentoxide (P2O5) per cubic meter of seawater (e.g., Atkins, 1923). Ironically, the chemical form P2O5 decomposes in water; the correct form should be P4O10 (Olson, 1967). Despite a logical recommendation by Atkins (1925) "to convert the conventional P2O5 values into the more rational values for the phosphate ion the factor 1.338, or very approximately 4/3, may be used to multiply the former," the P2O5 equivalence reporting practice continued. In 1933, Cooper (1933) made another plea for the importance of consistency in reporting dissolved nutrient and other elemental data. He suggested the gram-atom (or submultiple thereof, e.g., milligram-atom, microgram-atom) of the element under investigation per cubic meter as the most useful and meaningful concentration unit. This would provide for the direct comparison with other elements, and a relatively straightforward calculation of bioelemental atomic stoichiometry (i.e., C:N:P:Si) for dissolved or particulate matter. Atomic, molecular and ionic ratios would all be numerically identical. Cooper (1933) went on to state, "it is felt that such a radical change in the method of reporting results, before being put into service, requires the concurrence of the majority of oceanographical chemists, as uniformity in practice above all else is desirable." This bold suggestion was not

258

Karl and Bjorkman

immediately accepted by the contemporary community of scholars, and even at the present time there is no uniformity of reporting dissolved and particulate matter P concentrations. A variety of units, all interchangeable, have been used to report DOP in seawater. In preparing this review we have converted all of the reported concentration data to either ng-at P L~^ (nM P) or /xg-at P L~^ (/xM P) as appropriate. For organic P pools this refers to P only; so 507 ng L~^ of dissolved adenosine 5'-triphosphate (ATP), for example, would be equal to 1 nM ATP, but 3 nM P because each mole of ATP contains three P atoms. This practice of reporting DOP in P molar equivalents is absolutely necessary because the exact chemical composition remains largely unknown. For quantitative measurements of polymeric compounds such as DNA, RNA, and lipid-P we also report the assumptions that we made regarding the mole percentage of P in the specific polymeric compound. Sometimes molality (mol kg~^ of seawater) rather than molarity (mol L~^ of seawater) is used so that one does not have to calculate changes in volume that occur due to variations in temperature, pressure, or salinity but, for the purposes of this review, we will consider these changes to be negligible.

III. THE EARLY YEARS OF PELAGIC MARINE P-CYCLE RESEARCH (1884-1955) Several pathfinding scientific studies, especially those conducted during the first half of the 20th century, provided a sound foundation for contemporary investigations of the marine P cycle. The creation of the Marine Biological Association of the United Kingdom in 1884 and dedication of their marine laboratory at Plymouth in 1888, and the creation of the Marine Biological Laboratory at Woods Hole, Massachusetts, in 1888 are especially noteworthy because of the major impact these two research centers have had, and continue to have, on the field of marine ecology and biogeochemistry. In 1903, working out of the Plymouth laboratory, Donald J. Matthews began a systematic study of the oceanographic features of the English Channel. His time-series research program that was later continued by Atkins, Cooper, and Harvey, led to a comprehensive understanding of the fundamental links between nutrients, phytoplankton, and fish production in the sea. Matthews (1916, 1917) is also credited with making the first reliable estimations of phosphate in seawater, and with the discovery of oceanic DOP. The colorimetric method that he selected was based on the Pouget and Chouchak reagent (sodium molybdate/strychnine sulfate/nitric acid) which yielded a yellow colored product the intensity of which was proportional to the amount of phosphate in the water sample. Because this method was not very sensitive, Matthews (1917) first had to concentrate the dissolved phosphate by coprecipitation using either anmionia or a mixture of ammonia and an iron salt. The former, a predecessor to the modem

Dynamics ofDOP

259

"MAGIC" technique (Karl and Tien, 1992), removed phosphate by adsorption onto magnesium hydroxide, Mg(OH)2, and the latter as ferric phosphate and ferric hydroxide. Using this laborious but robust method for samples collected near Knap Buoy in the EngUsh Channel, Matthews (1916,1917) made two very important observations: (1) the concentration of phosphate in seawater was approximately 0.85 /xM in December 1915, decreasing systematically to a minimum of 80% of total P) is evident. The most notable seasonal change in the P inventory is the shift from Fi dominance in winter to DOP dominance in late spring-summer. Redrawn with permission from Harvey (1955).

(1937) and Cooper (1938) also accelerated during the 1930s, leading, eventually, to an ecumenical theory of nutrient dynamics in the sea. By the early 1940s, the fundamental role of DOP in the marine P cycle was firmly established (Atkins, 1930;Kreps, 1934; Cooper, 1938;Redfield^^a/., 1937; Newcombe and Brust, 1940). The sustained time series investigations of the EngHsh Channel provided evidence for a seasonally variable pool of DOP, which at the height of the phytoplankton bloom in late spring to early summer was maximal (Harvey, 1950; Armstrong and Harvey, 1950; Fig. 4). Field studies conducted in the epipelagic waters of the Gulf of Maine (Redfield et al, 1937) and in Chesapeake Bay (Newcombe and Brust, 1940) revealed similar results. Confirmation of the presence of a significant pool of DOP in seawaters from diverse habitats further stimulated research to ascertain the sources and sinks of

262

Karl and Bjorkman

these potentially diverse, but biologically, relevant, compound classes. Until this time, bacteria had been considered to be the principal agents of DOP remineralization to P/, the preferred substrate for phytoplankton growth. However, careful laboratory experiments conducted by Chu (1946) documented the ability of selected bacteria-free phytoplankton cultures to utilize DOP, thereby providing a novel, alternative pathway in the marine P cycle. Presumably these microorganisms would be selected for during the sunmier months when P/ concentrations were low and DOP/P/ ratios were high which could promote a seasonal succession of phytoplankton species in certain habitats. It seems appropriate to end this section on "The early years of marine P-cycle research" with the pubhcation of Harvey's seminal monograph "The Chemistry and Fertility of Sea Waters" (Harvey, 1955). While his field observations concentrated mainly on the English Channel, the conceptual framework presented in this now classic volume received worldwide attention and provided the incentive for a large portion of the DOP research which followed during the next half-century.

IV. THE PELAGIC MARINE P CYCLE: KEY POOLS AND PROCESSES Compared to the more complex cycles of C, N, and S that are characterized and sustained by redox transformations, the marine P cycle appears rather simple (Fig. 1). With few exceptions, P in the sea is present in the pentavalent state (-1-5) as P04^~, whether as free orthophosphate or as P incorporated into either phosphate ester or phosphonate compounds. The presence of phosphite (P03^~) and phosphine gas (PH3) has been reported in selected anoxic marine habitats where they were formed and, at least in the case of POj^", consumed as part of the marine P cycle (Devai et ai, 1988; Gassmann, 1994; Schink and Friedrich, 2000). These reduced P/ derivatives are not likely to be formed in open ocean habitats. Despite this redox simplicity, P/ is rapidly assimilated to form a diverse spectrum of organic and inorganic derivatives. These compounds have key structural and metabolic functions and, therefore, are continually produced by all living organisms. Cellular P metabolism in the marine environment is complex. The transfer of phosphoryl groups is a fundamental characteristic of intermediary metabolism and is, therefore, crucial for life. Numerous enzymes share the ability to catalyze phosphoryl group transfer including phosphatases, phosphokinases, phosphomutases, nucleotidases, nucleases, phosphodiesterases, phospholipases, and nucleotidyl transferases and cyclases (Knowles, 1980). Of these enzyme classes, the phosphatases (mono- and diesterases), nucleotidases and nucleases have been most frequently studied in the marine environment (Fig. 5). The depolymerization reactions converting high-molecular-weight (HMW) DOP to intermediate- and

Dynamics ofDOP exudation, grazing, death, autolysis, virallysis

263

^ ^

Living Blomass

Pyrophosphatast

PMEase NTPase

3' and 5' Nucleotides

Pyro-P/

C-O-P monomers

t

f

NTPase

nuclease POEase

L

; RNA : : and DNA •

;—I

C-P monomers

1

t

J

PPase

,

C-P tyg^e

C-P Lyase

PMEase

1 .

I

• G-o-P : : polymers ;

i inorganic i I poiy-P/ ;

i

L

c-p

i

I polymers •

DOP POOL

I

hydrolysis, fragmentation, disaggregation

Non-living Particulate Matter

Figure 5 Schematic presentation of the role of selected enzymes (both dissolved and cell/organismassociated) in DOP pool dynamics. The production of detrital P, including both particulate and IMW/HMW dissolved inorganic and organic P pools provides key substrates (open boxes) for the specific enzymes (shaded boxes). The continued supply of monomeric compounds ( 10), Prochlorococcus would be selected for and thus would begin to dominate the photoautotroph assemblage. Karl et al (2001a) have recently hypothesized that the NPSG has selected for Prochlorococcus diwd against eukaryotic algae over the past several decades and the attendant domain shift has caused significant changes in ecosystem structure and nutrient dynamics including P-cycle processes. The stoichiometry of the TDN:TDP pools in the upper 0-100 m of the water colunm (mostly dominated by DON and DOP) also displayed variability on monthly, seasonal, and interannual time scales. For example, monthly observations revealed occasional high frequency changes in the N:P ratio of the total dissolved matter pool between consecutive cruises (e.g., during spring 1989 and spring 1997; Fig. 21). These features were characterized by decreases in the N:P ratio from values that were significantly higher than the Redfield ratio to values approximating it (Fig. 21). When observed, these events always coincided with pulsed inputs of inorganic nutrients as detected by elevated nitrate plus nitrite (N+N) inventories (Karl et al, 2001b). In certain years (e.g., 1991, 1993, 1996) these nutrient injections were either absent or, more likely, missed by the relatively coarse monthly frequency of our sampling program. We also observed several time periods of sustained, systematic change in the N:P ratio; e.g., January 1991 to July 1992, where

302

Karl and Bjorkman 30|

I I I I I I I I I I I I I I I I I I I

OJAJOJAJOJAJOJAJOJAJOJAJOJAJOJAJOJAJOJAJOJAJO 1989 1990 1991 1992 1993 1994 1995 1996 1997 1998 1999 Sampling Date

'

DU

' 1

w

I'llllll III III

OJ

?)30

0)

n O20

llllllllll , ..ill

10 n

15

10

1 J

0

|20 Q.

Q ^-18 17

u i1 1

Spring Summer

1

i

1

20

25

-

I.1I

30

35

TDN:TDP (mohmol)

1

^21

1

^

III I I I lllllll III

fe40

E

1

III

1

50

^

- T

^

1

Fall

Season

Winter

1991 1993 1995 1997 19 Year

Figure 21 Nitrogen-to-phosphorus (N:P) ratios for the total dissolved matter pools in the upper 0-100 m of the water column at Sta. ALOHA during the 9-year observation period. (A) TDN:TDP versus sampling date. For each cruise, the mean value ±1 SD is presented. As a point for reference, the horizontal dashed hne is the Redfield ratio of 16N:1P. (B) Frequency histogram of TDNiTDP values for the 9-year data set. As a point for reference, the vertical dashed line is the Redfield ratio of 16N: IP. (C) Seasonal variability in TDNiTDP at Sta. ALOHA. Spring, Mar-May; Sunmier, June-August; Fall, September-November; Winter, December-February. The values presented are the mean d=l SD for each data set. (D) Interannual variability in TDN:TDP at Sta. ALOHA. The values presented are die mean ±1 SD for each data set. Redrawn from Karl et al. (2001b).

Dynamics ofDOP

303

the N:P decreased from >20 to values equivalent to the Redfield ratio, followed by an approximately 18-month period during which the N:P ratio slowly increased back to values approaching 25:1 (Fig. 21 A). These features resulted in significant interannual variations in the TDN:TDP ratio (Fig. 21D), with 1993 standing out as a year with an anomalously high mean TDNiTDP ratio of 22.8 (SD = 1.8). The mean TDN:TDP ratio for the complete 9-year data set was 19.6 (SD = 2.6), well above the 16N: IP Redfield ratio. Major differences were also observed for the molar N:P stoichiometrics of the dissolved inorganic nutrient pools (i.e., N+N:SRP) versus the total dissolved nutrient pools (i.e., TDNiTDP; Fig. 22). The greatest differences were observed in the upper 0-400 m of the water column (and, especially in the upper 0-100 m) where dissolved organic nutrients are present as significant fractions of the TDN and TDP pools. Whereas the dissolved inorganic N:P ratios in the upper water column were significantly lower than the Redfield ratio of 16N:1P, the N:P stoichiometry of the total dissolved pool (inorganic plus organic) was significantly greater than the Redfield ratio by as much as 50% (Fig. 22). Furthermore, there were systematic changes in the N:P stoichiometry as a function of water depth; inorganic N:P increased toward a ratio of approximately 14, while total N:P decreased toward the same value (Fig. 22). In both data sets, the greatest rate of change in N:P with depth was in the 100- to 400-m region of the water column. The relatively high TDN:TDP ratios in the near-surface waters are consistent with the hypothesis that P, not N, is the (or one of several) production rate limiting nutrient(s) in this ecosystem. This conclusion assumes that the TDN and TDP pools are fully bioavailable (see Smith, 1984; Jackson and Williams, 1985). However, recent research on dissolved organic matter suggests that near-surface pools are composed of at least two components: one that is locally produced and consumed during microbial metabolism (the labile pool), and one that may be more refractory. Although it is impossible to quantify these subcomponents using existing analytical techniques (and in reality there may be a continuum of bioavailabilities) for the sake of the present discussion we will assume that the mean deep-water (>600 m) DON and DOP pools are refractory. If TDN and TDP are corrected for these nonlabile components, the depth profile of N:P ratios assumes the characteristic "T-shape" (Fanning, 1992), but for a fundamentally different reason than the original author suggested. Rather than being a consequence of analytical uncertainties at low surface ocean concentrations, we hypothesize that the T-shaped profile for the corrected TDN:TDP ratios at Sta. ALOHA is a manifestation of an alternation between periods of N limitation (left-hand portion of the T) and periods of P limitation (right-hand portion of the T). Dinitrogen (N2) fixation is one of two major microbiological processes (the other being denitrification) that can significantly influence oceanic N:P stoichiometry on global scales. Several lines of evidence from Sta. ALOHA suggest that

Karl and Bjorkman

304 N:P (molimol)

N:P (mohmol)

N:P (mohmol) 10

20

0

5

10

15

20

25

30

•B 500

1000

Figure 22 Nitrogen-to-phosphorus (N:P) ratios versus water depth for samples collected at Sta. ALOHA during the period October 1988 to December 1997. (Left) Molar N:P ratios for dissolved inorganic pools calculated as nitrate plus nitrite (N-l-N):soluble reactive phosphorus (SRP). (Center) Molar N:P ratios for the "corrected" total dissolved matter pools (see text for details). (Right) Molar N:P ratios for total dissolved matter pools, including both inorganic and organic compounds, calculated as total dissolved nitrogen (TDN):total dissolved phosphorus (TDP). As a point for reference, the vertical dashed line in each graph is the Redfield molar ratio of 16N:1P. Redrawn from Karl et al. (2001b).

N2 fixation is an important contemporary source of new nitrogen for the pelagic ecosystem of the NPSG. In addition to the observed secular changes in SRP inventories and the N:P ratios already discussed, other independent measurements include: (1) Trichodesmium population abundances and estimates of their potential rates of biological N2fixation,(2) seasonal variations in the natural ^^N abundances of particulate matter exported to the deep sea and collected in bottom-moored sediment traps, and (3) increases in the DON pools during the period of increased rates of N2 fixation (Karl et aU 1997).

Dynamics ofDOP

305

At the beginning of the HOT program in 1988, biogeochemical processes in the gyre were thought to be well understood. New and export production were limited by the supply of nitrate from below the euphotic zone, and rates of primary production were thought to be largely supported by locally regenerated nitrogen. The contemporary view recognizes the gyre as a very different ecosystem (Karl, 1999; Karl et al, 2001a). Based on decade-long data sets, we hypothesize that there has been a fundamental shift from N limitation to P limitation (Karl et al, 1995; Karl and Tien, 1997). The ecological consequences of this hypothesized N2 fixation-forced P/ limitation, especially on DOP pool dynamics, is presented elsewhere (Karl et al, 2001b; see also Fig. 23). Suffice it to say that enhanced P/ cycling rates, shifts in the chemical composition of the DOP pool, and microbial biodiversity changes are all relevant features of these decade-scale ecosystem processes. The fundamental role of nutrient dynamics in biogeochemical processes

1970'S

N/P = 16

t

N/P =

16

1965

i

N/P =

16

"ENSO DECADE" (1982/3,1986/7.1991/3)

N/P = 40

I i

N/P = 16

Time N-limited Trichodesmium absent

N/P = 25

1980

N/P = 25

Time

1995

P-limited Trichodesmium present

Figure 23 Schematic presentation of the NPSG alternating ecosystem state hypothesis. This cartoon depicts the contrasting N and P nutrient cycles during periods of low rates of N2 fixation (e.g., 1970s) and enhanced rates of N2 fixation (1980-present). It is believed that the increased frequency and duration of the El Nino-Southern Oscillation (ENSO) cycle since the early 1980s is a major cause of the N2 fixation rate enhancement (see Karl, 1999; Karl et al, 2001a). The small rectangles and ovals at the top of each panel represent the average N:P ratios in particulate and dissolved matter, respectively, and the upward and downward arrows are the N:P stoichiometry of imported (mostly dissolved) and exported (mostly particulate) matter. N2 fixation (on right) decouples the N;P stoichiometry of the NPSG ecosystem. The center panels depict the inventories of SRP during both phases of the cycle showing a secular decrease in SRP following the selection and growth of N2-fixing microorganisms, such as Trichodesmium. Many of these predictions have been confirmed during the 12-year study at Sta. ALOHA (see text).

306

Karl and Bjorkman

and ecosystem modeling demands that we have a comprehensive, mechanistic understanding of inventories and fluxes. Although the present ongoing ocean timeseries study at Sta. ALOHA has certainly not resolved all of these important matters, it does provide an unprecedented data set to begin the next phase of hypothesis testing.

VIII. DOP POOL CHARACTERIZATION A major analytical challenge in DOP pool characterization is the detection of individual compounds typically present at pM to nM concentrations dissolved in seawater medium containing approximately 35 g L"^ of inorganic salts. Preconcentration and separation using ion exchange resins, ion exclusion or similar chromatographic procedures or even lyophiHzation that have proven useful for the characterization of DOP in soil extracts and freshwater habitats (e.g., Minear, 1972; Hino, 1989; Nanny et ai, 1995; Espinosa et al, 1999) are generally not applicable for the analysis of marine DOP. Because abiotic synthesis of organic P is not likely to occur in the marine environment, both the presence of a detectable DOP pool, as well as its molecular weight spectrum and chemical composition are dependent upon biological, mostly microbiological, processes. If marine DOP is derived from living organisms, as it ultimately must be, then the molecular spectrum of P in living cells or in marine particulate matter should be a first-order inventory of DOP sources. The macromolecular composition (by weight percent) of an "average" bacterial cell is as follows: protein, 52%; polysaccharide, 17%; RNA, 16%; lipid, 9.4%; DNA, 3.2%; other, 100-fold in periods of minutes to hours, have been observed during shifts from nutrient-sufficient to minimal growth media (Ault-Riche et al, 1998). Cells deficient in poly-P/ are noncompetitive during periods of nutritional stress, whether acute or chronic (Komberg et al, 1999). For growth in a fluctuating nutrient environment, rapid uptake and storage of P/ would be a key survival strategy. There is also an intracellular transient accumulation of poly-P/ at the onset of Vi depletion which appears to be under control of the Pho regulon discussed below. This process, termed "poly-P/ overplus" (Voelz et al, 1966), is fundamentally distinct from "luxury uptake" of Vi which also results in poly-P/ formation and storage but does not require P/ depletion. Of the two processes, the poly-P/ overplus phenomenon is probably most important in the marine environment and especially so in the open ocean. For example, if near surface ocean microbes are exposed to alternating periods of high and low P/, as they probably are (Karl, 1999), then this could lead to a poly-P/ overplus response and intracellular sequestration of P as poly-P/. Consequently, the current view that poly-P/ would not be expected to exist in low nutrient open ocean seawaters may be terribly incorrect; a focused research effort on this topic should be undertaken.

E BIOLOGICALLY AVAILABLE P Regardless of the rigor and precision with which P-containing compound pools are measured, the ecological significance of these analytical determinations will be incomplete until reliable estimates of the BAP pool are routinely available. In addition to P/, which is generally the preferred substrate for microorganisms, the P contained in a variety of polymeric inorganic compounds, in monomeric and polymeric organic compounds and in selected P containing minerals is available to some or all microorganisms; indeed some microorganisms may prefer esterlinked P sources to free orthophosphate (Tarapchak and Moll, 1990; Cotner and Wetzel, 1992). However, the bioavailability of most organic-P pools depends on ambient P/ pool concentrations and on the expression of specific enzymes that control transport, salvage, and substrate hydrolysis. Because many of these enzymes are induced by low P/, bioavailability may be a variable, time- and habitat condition-dependent parameter, rather than an easily predicted or measured metric. Assessment of the BAP may also depend on the time scale of consideration; for example, substrates that appear to be recalcitrant on short time scales

330

Karl and Bjorkman

(e.g., 50%) by the smallest size fraction (0.2-1 /xm), presumably auto- and heterotrophic bacteria. Fi turnover times were rapid, less than a few hours for most of the 3-month observation period.

338

Karl and Bjorkman

Release of incorporated ^^P from various particulate size-fractions was investigated by incubating with ^^P/ for a 3-h period, followed by the addition of an excess of unlabeled AMP (100 /xM). The addition of AMP, they reasoned, would partially inhibit the assimilation of recently produced [^^P]DOP compound and provide a more accurate estimate of gross DOP fluxes. The measured rates of [^^P]DOP release in these experiments were low, generally < 1 % of the corresponding particulate ^^P activity per hour (Dolan et al, 1995). However, when the concentration of oligotrich ciliates (predators of microorganisms in the 1- to 6-/xm size class) was artificially increased, there was a significant transfer of ^^P from particulate to dissolved pools. These field results confirmed the role of protozoan grazing in nutrient cycling processes (Andersen etal, 1986), including the P/ -^ PP ^- DOP -^ P/ pathway. Thingstad et al (1993) conducted a comprehensive study of microbial transformations of P in P-limited Sandsfjord, western Norway, including the coupling of P/ uptake, DOP production, specific DOP compound hydrolysis, and enzymatic hydrolysis. They focused on the production and turnover of nucleotides, and used ATP as a "model" compound. DOP/P/ concentration ratios in this habitat varied considerably but were generally between 10 and 100:1; P/ (reported as SRP) was 99% of total DOP pool) by polymeric compounds (presumably RNA and DNA) that turned over very slowly compared to the relatively small but rapidly assimilated nucleotide pool. In this regard, their results are consistent with the nuclease/phosphodiesterase "bottleneck" hypothesis discussed in a previous section of this review (Fig. 5). Orrett and Karl (1987) reported 0-100 m depth-integrated DOP production rates ranging from 0.3-0.8 mmol P m"'^ day"^ (TDP specific activity model) and 0.5-1.2 mmol P m"^ day~^ (RNA-specific activity model) for water samples collected in the NPSG. They reasoned that these DOP production rates could be further extrapolated to organic carbon, if the mean C:P molar ratio was either known or correctly assumed. An upper bound on C:P was taken as the whole cell C:P (106C: IP; Redfield et al, 1963), although it is unlikely that DOP compounds are, on average, this carbon-rich (see Table I). They assigned a value of 3C:1P as the theoretical lower bound on this value, a value identical to the nucleotide triphosphate pool. It is equally unlikely that the DOP pool would be that C poor, relative to P. A ratio of 9.5C: IP, the approximate value for RNA, was taken as the most reasonable estimate; the true C:P ratio for DOP is likely to be closer to the lower bound than to the upper bound. The extrapolated rate of DOC production, 24 mmol C m"-^ day~^ was about 50% of net primary production for this region (Karl et al, 1996,1998). Because this estimate is based on accumulation (net production) of DOP during the incubation period and, therefore, does not include

Dynamics ofDOP

339

contemporaneous [^^P]DOP production and [^^P]DOP utilization, gross DOP fluxes will be even larger. These results suggest an important role for DOP in microbial loop processes in these low-nutrient, open-ocean habitats. At steadystate, DOP production and DOP remineralization rates would be in balance. Consequently, given the DOP pool size and estimates of DOP turnover rates, these organic pools are likely to serve as an important source of P, as well as C and N, for microorganisms. If the compound C:P ratio is less than the whole cell C:P ratio, or if C is derived from additional or alternative sources, then P/ is likely to be released into the medium. These coupled processes most likely sustain the marine P cycle in the euphotic zone of the sea. Finally, Bjorkman et al (2000) measured coupled rates of P/ uptake and DOP production at several stations in the NPSG using ^^P/ as a tracer. Vi uptake rates varied from 3 to 8.2 nM Vi day"^; P/ pool turnover time was 2-40 days. Net DOP production (i.e., accumulation) was 10-40% of the net P/ uptake. The estimated turnover time for the entire DOP pool, assuming compositional singularity with nascent DOP, was 60-300 days. In all likelihood, the recently produced materials are assimilated much more rapidly than this simple calculation would suggest. Vi regeneration from selected, exogenously added DOP compounds was rapid and efficient; highest rates of P/ release were observed for nucleotides (Bjorkman ^r a/., 2000). Although coupled P/ uptake and DOP production is well documented in a variety of marine ecosystems, the actual mechanisms of DOP production remain elusive. Admiraal and Werner (1983) investigated the production of DOP by two coastal marine diatom species in laboratory culture. In addition to total DOP production rates, they concentrated a fraction of the DOP pool, using Sephadex G-10 chromatography, and documented partial reabsorption of the isolated DOP compounds by the same two species during P/-limited growth. The inadvertent diffusive loss of LMW compounds or the active excretion of both LMW and HMW compounds are both feasible; the list of specific compounds that are liberated by growing algae is very large (Fogg, 1966; Hellebust, 1974). Alternatively, DOP release could result from cell autolysis, predator grazing or viral lysis. Each separate pathway might be expected to produce a different spectrum of compounds. Most of the research conducted to date has focused on extracellular production of DOC, not DOP, but suffice it to say that DOP production by healthy microorganisms is probably a universal phenomenon.

B. DIRECT UTILIZATION OF DOP DOP compounds in seawater consist of both labile and refractory compounds. The labile DOP pool includes both transportable and nontransportable organic compounds, either of which can serve as P sources for microbial growth. The

340

Karl and Bjorkman

outer membrane of Gram-negative bacteria allows the transport of molecules up to about 600 Da (Weiss et ai, 1991). Therefore, many DOP compounds can, and probably are, taken up directly without the need for prior hydrolytic alteration. For example, Gly-3-P and AMP can be assimilated intact by certain bacteria (Wanner, 1993; Ruby et ai, 1985), whereas larger DOP compounds must be enzymatically hydrolyzed, either at the cell surface (or in the periplasmic space for bacteria) or in the surrounding medium prior to assimilation. The Pi released is then available for assimilation and biosynthesis. Therefore, the ability of an organism to grow on one or multiple DOP substrates as the sole source of cellular P can be traced to one of two independent properties: the presence of cell membrane or periplasmic bound enzymes that catalyze the DOP compound dephosphorylation and thereby enhance P/ availability or the presence of a DOP compound- or compound-class-specific uptake system. Both pathways are present in marine microorganisms; growth of both prokaryotic and eukaryotic microorganisms on a variety of different DOP compounds is well documented (Kuenzler, 1965; Cembella etai, 1984a,b; Antia etaL, 1991; van Boekel, 1991). Among other functions, the Pho regulon controls the transport of selected, intact DOP compounds into bacterial cells. Several proteins of the outer membrane of many bacteria (termed "porins") are involved in the formation of aqueous pores through which small hydrophilic molecules ( 1 /xm) particulate matter. The upper size boundary for colloidal matter lies at the juncture where gravity becomes the dominant force acting upon the particle. In essence then, a traditional view of colloids is that they are particles (not dissolved solutes) that do not sink unless they become entangled with other colloidal particles or sorb to sinking particulates. A 1-nm spherical diameter roughly equates with macromolecules of ^^1000 nominal molecular weight (or 1 kDa), a size equivalent to fulvic acids and marine porewater organic macromolecules (Chin and Gschwend, 1992). Organic biogeochemists traditionally categorize these and larger organic macromolecules as high-molecular-weight matter and characterize it in terms of elemental and molecular composition rather than its bulk interface characteristics. As a consequence, studies of the cycling of high-molecular-weight matter focus largely on specific molecular or biologically mediated chemical transformations. Casting the veil of "colloid" over macromolecular constituents does not diminish the importance of these processes but simply adds to them a range of nonspecific surface interactions that also might influence their behavior. In fact, most of the current dispute over the immediate fate of colloidal/high-molecular-weight organic matter lies in whether its short-term behavior is dominated by specific, biologically mediated chemical transformations or by rapid (and likely) nonspecific aggregative processes. By classical definition, the marine colloidal phase encompasses heterotrophic and phototrophic bacteria; termed "biocolloids." However, most oceanographers are dissatisfied with the concept of "biocolloids" so in most cases seawaters are filtered (0.2-0.8 /xm) to arbitrarily separate matter into a "particulate" phase, containing cells and large detritus, and a "dissolved" phase, containing solutes and colloidal particles. Operationally then, marine colloids are a subset of the classical colloid fraction described above. It has been argued recently that the definition of marine colloids instead should be based upon physicochemistry of the intracoUoidal matrix rather than a strict physical dimension (Gustafsson and Gschwend, 1997). In this "chemcentric" approach, the term colloid is applied only to those constituents that provide a molecular environment for the selective escape of chemicals from aqueous solution, by

371

Marine Colloids and Trace Metals

Gravitoidal

0.001 diameter (p.m) Figure 1 A graphical depiction of a chemcentric definition for colloidal matter. Here, the effect of particle dynamic processes on the mass distribution over the various size classes is illustrated. Steady-state particle size distributions (mass-based) are shown for three aquatic regimes having large differences in solids concentrations. The inflection point of each line where the slope changes is the functional distinction between colloids and "gravitoids" (sinking particles). The shaded areas depict the traditional, operational boundaries between soluble, colloidal, and "particulate" fractions. Based on a chemicentric approach, the upper size boundary of colloidal matter shifts in conjunction with the total solids concentrations in the water. Reprinted with permission from Gustafsson and Gschwend (1997).

either partitioning into or onto the colloidal constituent. By this definition, the lower threshold separating solutes from colloids still corresponds to a physical dimension of ~ 1 nm (for the reasons outlined above) but the upper size threshold is constrained by environmental transport conditions rather than by arbitrary size delimitation (Fig. 1). In a refinement of the classical colloid definition, the size boundary between colloids and "gravitoids" is determined by the outcome of kinetic competition between aggregation and sedimentation. Gustafsson and Gschewnd (1997) argue that this size boundary shifts as a function of the total solids concentration, so the upper threshold delimiting colloidal matter will be several micrometers in coastal waters versus several tenths of microns in the deep ocean (Fig. 1). Another significant distinction is that not all substances larger than a nanometer are colloidal. High-molecular-weight polyelectrolye molecules that assume an extended conformation in seawater would not meet the chemcentric criteria for colloids. This aspect is problematic for studying the marine colloidal phase because no methodologies currently exist for measuring the conformation

372

Mark L Wells

of organic molecules in marine water samples. However, a functionally based definition is more adaptable to studying the effect of colloidal processes in natural systems (Gustafsson and Gschwend, 1997). While the past decade has brought more dissention than consensus about what constitutes marine colloids, our understanding of the varied roles that colloids may play in coastal and offshore waters has nonetheless improved despite arbitrary and inconsistent size-based delimitations of the marine colloidal phase. The analytical methods underlying these studies are now considered.

III. ANALYTICAL METHODS The analytical approaches used to study marine colloids lie in two broad categories; determination of the abundance and elemental and molecular composition of marine colloids, and the assessment of colloid reaction rates, primarily with respect to their transfer into particulate phases. The methods used to quantify the abundance of colloidal matter will be considered now while the measurement and implications of colloid reaction rates are covered in section VII.

A. NUMBER CONCENTRATIONS OF MARINE COLLOIDAL MATTER Early studies established that a significant fraction of dissolved organic matter lay in the colloidal size range (e.g., Carlson et al, 1985; Maurer, 1976; Moran and Moore, 1989; Ogura, 1977; Sharp, 1973). These findings catalyzed a burst of interest in marine colloids and their role in carbon and metal cycling. Koike et al. (1990) reported that the abundance of nonliving organic particles sized between 0.38 and 1.0 /xm were ~10^ particles mL~Mn surface waters of the North Pacific. This finding was corroborated for coastal waters off Nova Scotia in a joint project using several different analytical approaches (Longhurst et al, 1992). The concentration of these "Koike" particles decreased by 10x in deep waters, implying that there was active production of colloidal matter in the photic zone. These flexible (difficult to filter) particles were 4-30 x more abundant than marine bacteria. Moreover, Koike et al. (1990) showed that bacteria were not a source of these colloids but that they likely originated with the activity of small flagellates. They estimated that these particles accounted for ~10% of the DOC, in agreement with estimates from earlier bulk separation studies (Sharp, 1973). Number concentrations of marine colloids in seawater were measured for sizes down to ^ 5 nm using a combination of ultracentrifugation and transmission electron microscopy (Wells and Goldberg, 1991, 1992, 1994). Colloid concentrations

Marine Colloids and Trace Metals

373

increased logarithmically with decreasing size in coastal California seawaters, with numbers being on the order of 10^ colloids niL~^ Similar abundances were observed in surface and deep waters of the North Atlantic, equatorial Pacific, and the Southern Ocean (Wells and Goldberg, 1994, and unpubHshed data), demonstrating the widespread distribution of marine colloidal matter. At each station, the mean particle size tended to be larger near the base of the thermocline, suggesting a different source or intensity of colloid production in this region. In all cases, the globular-shaped colloidal particles exhibited heterogeneous electron densities, suggesting that they perhaps are aggregates of smaller molecules (Fig. 2). In subsequent studies, resin embedding methods were used to ensure that molecules maintained their configurations upon drying, and stains were applied to improve the visibiHty of the colloidal phase (e.g., Heissenberger and Hemdl, 1994; Leppard et al, 1997). These studies confirmed the presence of the granular colloids noted above and showed that additional, more amorphous colloidal matter also was present. Transmission electron microscopy (TEM) studies also showed the presence of large aggregates of colloidal matter (Leppard et al, 1997; Wells and Goldberg, 1993) that can become incorporated into marine snow aggregates (Leppard ^f fl/., 1996). More recently, Santschi et al (1998) used a combination of TEM and atomic force microscopy to show that fibrillar colloids, 1-3 nm in width and 100-2000 nm in length, comprise a significant fraction of colloidal organic matter in coastal and offshore seawaters. Thesefibrils,rich in polysaccharides (Santschi et al, 1998), are clearly colloidal by the standard size definition but may be noncolloidal based upon a chemcentric view (Gustafsson and Gschwend, 1997). Regardless, these fibrils form aggregates up to several micrometers in size, often incorporating globularshaped colloids (Santschi et al, 1998). Fibrillar "particles" therefore are likely to be important in colloid cycling. Dynamic light scattering (DLS), also known as photon correlation spectroscopy, has been successfully applied recently to the study of colloidal abundance and formation in seawater (Chin et al, 1998). With dynamic Hght scattering, the time dependence offightscattered from a laser-illuminated volume of solution is measured over tenths of a microsecond to milliseconds. These fluctuations are a function of the diffusion rate of molecules and particles within this volume (that is, Brownian motion). The time dependence of scatter therefore can be used to calculate the diffusion coefficient of particles if a number of conditions can be met. In favorable cases there are methods available for treating the time-dependent fluctuations in the scattered light intensity to extract the "hydrodynamic" (or "Stokes radius") colloid size distribution. Chin et al (1998) used this capabiHty to measure shortterm changes in the abundance of colloids a few nanometers to a few micrometers in size (see below). This methodology is certain to be applied more frequently in future studies on marine colloid dynamics.

374

Mark L. Wells

10m

300 m

50 m

2500 m

100 m Figure 2 Transmission electron microscope images of marine colloidal matter from a depth profile in the Sargasso Sea. Colloids were settled directly onto charged TEM grids by ultracentrifugation (see Wells and Goldberg, 1994). Differences in colloid size, morphology, and abundance are readily apparent, with the larger colloids being most prevalent near the base of the photic zone (100 m) and immediately below (3(X) m). These globular colloidal particles as well as more electron-transparent colloidal particles are seen in nearshore waters when embedding and staining methods are employed (Leppard et ai, 1997). Scale bars, 100 nm.

Marine Colloids and Trace Metals

375

B. ISOLATION OF COLLOIDAL MATTER FOR BULK ANALYSIS Analysis of the constituents that comprise marine colloidal matter generally requires the separation and preconcentration of colloids from conventionally filtered (e.g., 0.2 /xm) seawater. Although it is possible to analyze elemental compositions of individual particles using TEM/energy dispersive spectroscopy (e.g., Chin etal, 1998), the low sensitivity of the method and nonhomogeneity within and among individual particles strictly limits the quantitative value of this approach (Wells and Goldberg, 1991). The primary method at present for preconcentrating colloidal matter for bulk chemical analysis is cross-flow (or tangential flow) filtration (CFF). This approach is attractive for its operational simplicity and the high concentration factors (>100x) that can be achieved. However, the separation of molecules by pore size exclusion is strongly influenced by molecular conformation, interaction with the membrane and interaction with other soluble and colloidal substances near the membrane surface (Buffle et al, 1992). Aside from conformational and molecular flexibility issues that cloud the accuracy of size exclusion methods, solvent flow through the membrane leads to the accumulation of macromolecular substances near the membrane surface; a process that is countered by back diffusion of molecules from the membrane. This concentration "polarization" can enhance colloid-colloid and colloid-solute interactions. Small solutes that should otherwise pass through the membrane might then become associated with larger colloids that do not, altering the apparent size fractionation. Directing sample flow tangentially across the membrane surface reduces the thickness of the concentration polarization layer, decreasing but likely not entirely eliminating the possibility for self-aggregation (Buffle et al, 1992). The lowering of the osmotic barrier also increases permeate (filtrate) flow rates. In practice, the retentate solution containing the macromolecules is swept from the membrane and recycled through the retentate reservoir. This reservoir usually encompasses the entire starting sample volume, but in a few cases a small retentate reservoir instead is continuously replenished with fresh sample water as CFF proceeds (e.g., Gustafsson et al, 1996). The latter approach, termed sampling mode (Dai et al, 1998), minimizes the exposure of the colloidal constituents to the CFF system and may yield better estimates of the retention coefficient of a molecule. Determination of the colloidal fraction of carbon or metals in conventionally filtered (0.2-0.7/xm) "dissolved" samples typically is done in one of two ways. The more straightforward method is to subtract analyte concentrations in the membrane permeate from those in the starting filtrate solution. But this approach potentially can bias the determinations because any sorption of truly soluble metals or organic molecules to the CFF system is then quantified as being colloidal, thus overestimating the colloid fraction. Conversely, the colloidal fraction could be underestimated if there is low-level carbon or metal contamination of the permeate

376

Mark L Wells

from the CFF system. Nonetheless, once system leaching and sorption problems are verified to be minimal for a given water type, ultrafiltration cartridges can offer a viable straightforward approach for determining colloidal metal concentrations (Nishioka^r^/., 2001) The preferred analytical approach is to determine mass balance for each sample separation to ascertain if there are contamination or sorption problems. In this case, analyte concentrations are measured in the starting filtrate, the membrane permeate, and the membrane retentate fractions. Colloidal metal concentrations are then calculated by subtracting the permeate concentration from the retentate and dividing the result by the concentration factor. Mass balance can then be assessed by comparing the sum of the analyte soluble and colloidal concentrations with that in the conventionally filtered starting solution. While preferable over the simple difference approach, mass balance determinations still are relatively insensitive and could mask significant sorption problems (Gustafsson et ai, 1996). A measured loss of analyte to the CFF system may not be due to sorption. Incomplete flushing while extracting the CFF retentate will leave a large portion of colloidal material in the concentration polarization layer, leading to a low estimate of the colloidal metal concentration (see in Buffle et al, 1992). For example, it is recommended that once CFF processing is complete, the retentate solution should be recirculated for some time with the permeate flow turned off to enhance recovery of the colloidal material (Buesseler et ai, 1996). In practice, any "missing" analyte usually is attributed to incomplete colloid recovery, the identical result as taking the simple difference between permeate and starting solution concentrations. Nonetheless, determining mass balance provides an indication of which results should be interpreted with added caution. The increasing use of CFF in colloid studies during the early 1990s led to an intercomparison study to assess whether different CFF systems provided well-defined and operationally reproducible results (Buesseler et al, 1996). This "colloid cookout" study was conducted using 14 different CFF systems representing five different manufacturers, with the central criterion being the size fractionation of organic carbon with 1-kDa membranes. Large volumes of homogenized surface waters off Woods Hole, Massachusetts, and mid-depth (600 m) waters off the National Energy Laboratory of Hawaii were processed on-site (Buesseler et al, 1996). Although the primary focus was testing the separation of colloidal organic matter, the outcome is sunmiarized here because it has direct significance to the study of colloidal trace metals. There were two primary findings of the intercomparison study. First, extremely long cleaning and flushing times are required to reduce the DOC blanks of new cartridges. Second, the degree of colloid retention by the 1-kDa membranes varied dramatically among manufacturers, but was similar among different groups using the same brand of membrane. Even so, retention efficiencies can vary among CFF membrane batches from a single manufacturer (Dai et al, 1998, P. Santschi,

Marine Colloids and Trace Metals

377

pers. commun.) reflecting the shortcomings of using industrial-based separation technologies outside the limits of their design application. For systems displaying high retention efficiencies of marine colloids (e.g., Amicon), permeate DOC concentrations increased significantly with the concentration factor. In contrast, this "breakthrough" of organic carbon was not apparent in systems equipped with membranes having lower colloidal retention efficiencies. A similar increase in the permeate concentration of trace metals has been observed with Amicon 1-kDa membranes (Guo et al, 2000b; Wen et al, 1996) but not with Filtron 1-kDa membranes (Powell et al, 1996; Wells et al, 2000). Although there is uncertainty about the cause of changing analyte concentrations in the permeate during processing (Buesseler et at., 1996), it may be due to increased membrane transport of soluble (40% of 0.5-kDa rhodamine 6G and 0.6-kDa glutathione are retained by the 1-kDa Amicon SlONl membrane, even at concentration factors of ^50. They suggested that the reverse problem, breakthrough of high-molecular-weight standards, was not significant (but see below). Retention of soluble (40) help to minimize the retention of soluble organic phases, contrary to earlier recommended protocols (e.g., Buesseler et al, 1996). They also showed that diafiltration with deionized water further reduced the retention of their < 1-kDa molecular probes. However, the mechanistic interpretations by Guo et al. (2000b) for the increasing permeate concentrations with higher concentration factors rely on a permeation model for single discrete molecules that assumes that the sorption of molecules to the membrane is neghgible (see in Kilduff and Weber, 1992; Logan and Juany, 1990). But sorption of certain substances can be significant (e.g., Dai et al, 1998; Gustafsson et al, 1996). This simplified model also may not apply equally well to complex mixtures of individual compounds or to colloidal assemblages of discrete molecules, both of which likely comprise the marine colloidal phase. The decreasing ionic strength during diafiltration of the retentate also may alter the conformation of natural colloidal organic molecules enough to affect their retention. For example, decreasing Mg^"^ and Ca^+ activities causes disaggregation of natural colloidal polymers in coastal waters (Chin et al, 1998). The question of the breakthrough of organic molecules during CFF is an issue of continuing debate. Dai et a/. (1998) compared the performance of the Amicon 1-kDa membrane used by Guo et al. (2000b) with the Millipore Prep/Scale 1-kDa membrane with standard molecules as well as nearshore and offshore seawaters. This comparison is particularly useful because the Amicon 1-kDa membrane, a mainstay for marine colloid studies, became no longer commercially available

378

Mark L.Wells

after Amicon merged with Millipore. From these data, Dai et al. (1998) concluded that breakthrough of both high- and low-molecular-weight matter occurs during processing as a consequence of the CFF membrane itself as well as physical/chemical interactions of specific organic constituents with the membrane. Breakthrough varied among the oceanographic sites likely due to differences in molecular composition and concentrations of COC. They attributed the bulk of this breakthrough to high-molecular-weight colloids, in contrast to the findings of Guo et al. (2000b), and reconmiended keeping CFP concentration factors 20 elements in stream waters. Unfortunately, the high salt content of seawater precludes this simple approach for all but perhaps estuarine waters, but interfacing flow-FFF with flow injection analytical methods might still provide a similar capability for studying colloidal trace metals in marine waters. While the very limited use offlowFFF in marine waters has prevented close examination of potential methodological artifacts influencing the size separations, these techniques are certain to gamer attention in the future. The differing and imprecise size retention of marine colloids by CFF membranes challenges our ability to quantitatively compare findings among studies from disparate locations and times. While this situation is unfortunate, it is important to remember that these shortcomings are not restricted to CFF systems. Conventional filtration of seawater also suffers major artifacts in size selectivity, even when using etched membrane filters (e.g., Koike et al, 1990; Stockner et al, 1989). Nonetheless, CFF presently remains the only effective method for accumulating enough high-molecular-weight substances from seawater to examine the broad molecular and metal composition of the marine colloidal phase. For that reason, CFF will continue to serve a key role in colloid studies in the foreseeable future.

IV. METAL CONTENT OF MARINE COLLOIDAL MATTER The tremendous increase in study of colloid-associated trace metals over the past decade has considerably expanded the database for estuarine and nearshore waters (Table I). Even so, this database comprises ^—50%o, but central North Pacific (A^'^C = -525 ± 20%^; 5980 years equivalent age). As we will describe below more fully, the vertical distributions of A^'^C-DOC in the openocean water column may be used in simple two-box mass balance models to show that DOC above the main thermocline may be described adequately by a combination of (a) ^"^C-depleted, subthermocline DOC that is mixed vertically over time scales of ocean water transport and cycling into the upper water column, and (b) ^^C-enriched DOC that is newly produced by recent biological production. Early work comparing the deep A^'^C-DOC profiles of the central North Pacific (WilUams and Druffel, 1987) with the Sargasso Sea (Bauer et al, 1992a; Druffel et al, 1992) emphasized the near-uniform offset in A^'^C (135%^) and age (2100 years) between these two profiles below the main thermocline. This age difference is close to the approximate 1500-year transport time of deep ocean water between the North Atlantic and North Pacific estimated by Stuiver et al. (1983). This led to the interpretation that deep-ocean DOC is refractory, it ages

421

Carbon Isotopic Composition of DOM

quasi-conservatively during global deep-water transport, and (because of the ages of DOC from both the Sargasso and North Pacific exceeding deep water transport times) it is recycled through the oceans several times prior to being removed. As we shall see, while this overall interpretation may still be valid, new data suggest that the details concerning the transport, aging and utilization of DOC through the deep global ocean may be more complicated than a simple comparison of the Sargasso and North Pacific "end-members" suggests. Although A^'^C-DOC values decrease along the path of mean deep water transport, they do not decrease at the same rate as A^^C of dissolved inorganic carbon (A^'^C-DOC; Fig. 4c), which is generally believed to be a robust measure of deep water mass aging during transport (Stuiver et al, 1983). As can be seen (Fig. 4a), the subthermochne A^'^C-DOC of the Southern Ocean is on average only sHghtly greater (-'25%o) than that of the central North Pacific. In contrast, the A^'^C-DOC in the Southern Ocean is essentially equidistant between the North Atlantic and the North Pacific values and profiles (Fig. 4c). Furthermore, Southern Ocean DOC is very similar in concentration to the Sargasso Sea (Fig. 4b). Therefore, while there is only about a 2 /xM decrease in DOC concentration between the North Atlantic and Southern oceans, the greatest part of the A^'^C-DOC decrease occurs in this same sector. The A^^C-based age differences for both DOC and DIC between these three oceanic regions (Table III) show that the age difference for DOC is about 900 years greater than for DIC in the Atlantic-Southern Ocean sector, whereas the two are much closer in the Southern Ocean-Pacific sector. Furthermore, assuming quasisteady-state rates of change of both A^'^C-DOC and DOC concentration as deep water ages and DOC is degraded, an estimate may be made of a "conservative" rate of change of both of these parameters between the Sargasso and North Pacific (Fig. 5). Taken together, these findings suggest that there may be (a) selective utilization or removal (~2 /xM) of ^"^C-enriched DOC between the Sargasso Sea and Southern Ocean, (b) a selective replacement of average DOC by ^^C-depleted DOC (~3 /xM of A ^^C = -1000%o material, such as petroleum or black carbon) in Southern Ocean waters, or (c) a combination of the two. There exists evidence both for the selective utilization of ^"^C-enriched DOC components by marine bacteria Table III Deep Water Transit Times and DOC Age Differences Sector

Transit time (years) (A^^C-DIC-based)"

DOC age difference (years)

Sargasso —> S. Ocean

700

1600

S. Ocean —> N. Pacific

600

400

^ Based on Stuiver et al. (1983)

422

James E. Bauer -350

-

Sargasso Sea A A ^ A

-400

-

Q

o

/

y

•

-550

1

32

^^f_i

34

X

^ *

A

Central N. Pacific

> ^

y

y

A

A

A ^

Southern Ocean A

^

• ^—

1

36

L.

1

38

r

_i

40

1^

1

42

1

1

44

1

46

DOC (|LiM) Figure 5 The relationship between A^"^C-DOC and DOC concentration in deep waters ( > ~ 1000 m depth) of the Sargasso Sea (open triangles). Southern Ocean (filled triangles), and central North Pacific (filled diamonds). The dashed curve represents the expected changes in A ^^C-DOC and DOC assuming a steady rate of change of each during deep water mass transport and aging. Adapted with permission from Druffel and Bauer (2000).

(Cherrier et aL, 1999) as well as for inputs of both ^^C-depleted petroleum carbon (Boehm and Requejo, 1986; Macdonald et al, 1993; Roberts and Carney, 1997) and black carbon (Masiello and Druffel, 1998) to open ocean waters. Whatever the case may be, it is clear from these and other studies (e.g., Hansell and Carlson, 1998) that deep ocean DOC is not completely refractory on time scales of deep water transport, and that both the utilization and input of components of different A^^C signatures may influence the distributions of concentrations and A^'^C of DOC throughout the deep ocean. There has been little work on the A^'^C and/or 5^^C signatures of different components of open ocean DOC. The only organic "components" (operationally defined or otherwise) that have so far been isolated from open-ocean DOC for A^'^C analysis are the humic substances (hydrophobic acids extractable on XAD resins) and the UDOC ( >1 kDa molecular weight) fraction (Table II). The range of A^^C-humic values for the central North Pacific is relatively narrow (—410 to —310%o) compared to that of the total DOC. Conversely, in the Sargasso Sea the range of A^'^C-humic values is greater (-587 to -358%o) compared to that of the total DOC. In both the central North Pacific and Sargasso Sea, A^^C-humic was significantly lower than A^'^C values for total DOC from the same depth (Bauer et ai, 1992a; Druffel et al, 1992). This indicates that this component of the DOC pool must be derived from older, possibly non-marine sources or from hydrophobic marine constituents such as lipids that are fighter in 6^^C (Sackett, 1989) or that it simply ages in situ more extensively than the average DOC.

Carbon Isotonic Composition of DOM

423

In the only known study of open-ocean UDOC in the major open ocean systems, and in contrast to humic substances, Bauer et al. (submitted for pubhcation) demonstrated that the > 1-kDa fraction of DOC (comprising ~40-50% of the total DOC in this study) had A^'^C and 8^^C values that were indistinguishable within analytical error to the total DOC (Table II). This suggests that, isotopically at least, the high-molecular-weight fraction of DOC is representative of the total pool in open-ocean settings. As we shall see, however, there is greater isotopic disparity between the bulk DOC and different molecular-weight components in more coastal settings (see Section IV.B, below). It should be noted that measurement of the isotopic composition of a bulk pool of carbon like DOC, the A^'^C or 8^^C signatures of that bulk pool reflect a weighted average of the isotopic signatures of all the components contributing to it. For example, considering the mean A^'^C-DOC values and corresponding ages for the deep Sargasso Sea (mean A^^^C = -390 ± 10%^; 3970 years B.P.) and central North Pacific (mean A^'^C = -525%^ ± 10%o; 5980 years B.P.), there are a number of possible scenarios for the A^'^C distributions of the entire population of organic molecules in the DOC pool (Figs. 6A-6C). Frequency distributions of A^'^C values of all the components of DOC may have varying degrees of normality as well as varying ranges. For example, frequency distributions of A^^C of different DOC fractions or even molecules may be broad and continuous (Fig. 6A), narrow and continuous (Fig. 6B), or even discrete (Fig. 6C). A number of other potential scenarios may be hypothesized as well. The specific combination of factors that leads to the observed mean weighted A^'^C-DOC values, and the distribution of A^'^C values and ages within the DOC pool is not known, though the advent of compound-class (Wang et al, 1998; Wang and Druffel, 2001) and compound-specific (Eglinton et al, 1996, 1997) A^^C measurements may allow for a greater degree of differentiation of the A^'^C distributions within the bulk DOC pool. The few compound-class A^^C measurements that have been made to date in oceanic DOC (isolated as UDOC) show that, for sugars at least, modem radiocarbon ages predominate, in spite of the old, ^^C-depleted DOC that dominates the average, bulk pool (Aluwihare, 1999). Thus, young, surface ocean-derived components of the DOC pool are transported to the deep ocean where they may fuel deep heterotrophic metabolism (Craig, 1971a,b; Nagata et al, 2000) or where they may become refractory and age (Brophy and Carlson, 1989; Ogawa et al, 2001). B. DISTRIBUTIONS OF ^^^C AND A ^ ^ C OF DOC IN O C E A N M A R G I N S

The distributions of S^^C and A^'^C of DOC and their use as source and age tracers, as well as their comparison to distributions in the open ocean, have only

424

James E. Bauer Deep Pacific

Deep Atlantic

73 O

•^ O

10

2 ^

Deep Pacific

^^,

Deep . ^ ^ ^ Atlantic

^s 0^ U

.>td: Q o 10 g

C/5

-O

O

-

5 ^

A

< ^ (D U

.>^: Q o 13 >::;

M

/\ \

ft, ii 1

10

Deep Pacific Deep Atlantic

(\

\W

«

I ,1 ;

5

A

i^CAGECkyrB.F.)

i^ 1

0

Figure 6 Three hypothetical distributions of ^ "^ C ages of constituent organic molecules within the bulk DOC pool. (A) Broad, continuous distribution; (B) narrow, continuous distribution; and (C) discrete distributions.

recently been established in selected ocean margin systems (Table IV). For purposes of this discussion, we will consider ocean margins and coastal environments to be those regions extending from the inner continental shelf, across the continental slope to the continental rise. A detailed description of 5^^C and A^'^C of DOC in areas farther inshore such as estuaries and rivers will not be undertaken here, except insofar as these systems may be important sources of DOC to shelf, slope, and rise waters. For a comprehensive review of 8^^C and A^'^C of DOC in river and estuarine systems, the reader is referred to Raymond and Bauer (2001a).

g

•53D

O

I u o

C8 01)

U

C/3

r--

'^

O

'^ O

in in J3 in in

i-K

^H O N r-H T-H O N

O N O N CN O N O N O N ON O N ON ON O N r^

^ O

,_i

(U

^ O

^ O

VH

^ O

u u^u u ^

s ^ G^ ^ ^

1-H

r—H

T-H

^

in 1 ON

1 1

i n CO CO CO

1 1 CN o o ^ rl I CO q q 1

^ So.\ A ; ^ ^ / ^ MAB ^ - - . ^ / J ??5 y^f^ slope J -A X^ A \ -??

J

30

^

„J

32

A

fc_

..i

34

.

^

j

J

36

Salinity (psu) Figure 8 (a) A^'^C-DOC vs ^^^C-DOC, (b) A^^C-DOC vs salinity, and (c) ^^^c-DOC vs salinity for shelf and slope waters of the Middle Atlantic Bight (MAB) region of the western North Atlantic in spring 1994. Also indicated in (b) and (c) are values at different depths in the Sargasso Sea (SS). For (a), the correlation between A^^^C-DOC and 5^^C-D0C yielded r = 0.706 (Pl-kDa fraction has a lower overall range than average bulk DOC measured in other studies (compare values for UDOC with total DOC in Table IV), indicating that it is composed of isotopically lighter material. The A^'^C of the >l-kDa fraction is generally similar to that of the total DOC from the same regions (Table IV), even though direct comparisons have not been made as in the open ocean (Table II, and see Section III.A, above). For >10-kDa UDOC, with the exception of the Mid-Atlantic Bight slope, 5^^C values were overall heavier, and A^^C values were correspondingly enriched, compared to either >l-kDa UDOC or total DOC in the Mid-Atlantic Bight and Gulf of Mexico (Santschi et al, 1995; Guo et al, 1996; Guo and Santschi, 1997; Table IV). One interpretation of these findings (Santschi et al, 1995; Guo et al, 1996) is that the higher molecular weight (> 10-kDa) components of DOC include a disproportionately large contribution from young, ^"^C-enriched marine organic matter compared to either the total or 1- tolO-kDa fractions, which themselves may contain a significant contribution from older, terrigenous inputs. On the very dynamic Mid-Atlantic Bight continental slope, however, the >10-kDa fraction is highly depleted in both A^'^C and ^^^C (Guo et al, 1996; Table IV). In fact, the most depleted A^'^C values in any form of DOC ever observed are found in this fraction in Mid-Atlantic Bight slope waters. Subsequent work by Guo and Santschi (2000) suggests that these anomalously low A^'^C and 5^^C slope values may result from preferential desorption of ^"^C- and ^^C-depleted organic matter during reworking and resuspension of slope sediments. Thus, old ^"^C-depleted DOC and UDOC from slope sediments and benthic nepheloid layers, and possibly including components derived from terrestrial sources, may contribute in part to the old DOC in the open ocean (Bauer and Druffel, 1998). Below we show how simple two- and three-source mass balance models have been applied to A^'^C-DOC and 5^^C-D0C distributions in ocean margins to evaluate (a) the relative contributions of known (in terms of their A^'^C and 5^^C signatures) potential sources of marine and terrestrial DOC to the standing inventory of shelf and slope DOC of selected margin systems and (b) the potential contributions of deep margin (i.e., slope) DOC to the deep open ocean.

430

James E. Bauer

V. APPLICATIONS OF 6^^C AND A^^C IN MARINE DOC CYCLING STUDIES Both (5^^C-D0C and A^'^C-DOC have been used in simple modeling studies of marine organic carbon cycling. The most common application of S^^C has been in evaluating the relative and/or absolute magnitudes of the contributions of different organic matter sources to a given pool or reservoir. 8^^C has been most successfully applied in coastal, estuarine, and other systems where the various potential inputs have measurably different signatures. As discussed above, in openocean water (Table II) the 5^^C of DOC has such a small range (and the factors that are responsible for the limited observed variability are not well-enough established) as to essentially negate its use as a source tracer in these systems. In coastal systems, where contributions of organic matter from nonmarine sources to the DOC pool may be more important and the range of 5^^C-D0C values is greater (Table IV), 8^^C may be used in a semiquantitative manner to establish the relative contributions of different sources. Furthermore, and in contrast to 5 ^^C, A ^"^C-DOC has a greater natural range in both oceanic (Table II) and coastal (Table IV) systems and thus may prove more useful as a potential "source" tracer than 8^^C. It should be noted that when we refer to both isotopes as source tracers, it is in the sense that 8^^C can differentiate between different sources on the basis of fractionation effects (i.e., between different primary producers), while A^^C can differentiate on the basis of source age. Used in conjunction with each other, A^^C and 8^^C may serve as unique source and age tracers in different environments (e.g., in oceanic vs coastal systems) and at times may be used to compliment one another as dual tracers of carbon specifically, and organic matter in general. Natural (cosmogenic) and bomb ^"^C have the added benefit of providing information on the mean weighted age of a given carbon pool. We will now consider some of the major applications of A^'^C and 8^^C as both source and age tracers of organic matter in studies of marine carbon cycling through the DOC pool. A. VERTICAL MIXING AND DISTRIBUTION OF A ^ ^ C - D O C IN THE O P E N O C E A N

Profiles of bulk A^'^C-DOC values and DOC concentrations in all three major open ocean regions studied to date (Figs. 4a and 4b) indicate that there are decreasing inputs to the open ocean water column of excess (i.e., above deep, background A^'^C-DOC values) ^"^C-enriched DOC with increasing depth. Furthermore, the A^'^C-DOC profiles suggest (a) that there must be inputs of the most ^^C-enriched (i.e., "young") material in the shallowest parts of the water column, (b) that these ^"^C-enriched inputs decrease with increasing depth through the main thermocline, and (c) that there are no discernible inputs in ^"^C-enriched DOC below the main thermocline due to the absence of vertical A^'^C-DOC gradients in deep waters

431

Carbon Isotopic Composition of DOM (although this observation alone is not sufficient to rule out such inputs to the deep ocean by other means such as POC dissolution and degradation, lateral inputs, porewater diffusion, etc.). The question, then, is how to interpret the A^'^C-DOC distributions and gradients throughout the open ocean water column. In order to address this, Williams and Druffel (1987), Bauer et al. (1992a), and Druffel et al (1992) invoked a two-component vertical mass balance mixing model that assumed (1) deep-ocean DOC, with its near-constant average concentration and A^'^C signature, is mixed homogeneously throughout both the deep and upper water columns; (2) surfaceocean DOC is composed of a combination of this ^"^C-depleted deep-ocean DOC and ^"^C-enriched DOC derived from contemporary surface-ocean productivity; and (3) the A^'^C values of the contemporary surface-derived DOC are equivalent to the A^'^C of surface ocean DIC from whence the DOC is fixed. Taking the Sargasso Sea as an example (Fig. 9A), and using the above assumptions, Bauer et al. (1992a), and Druffel et al. (1992) predicted that the average mixed layer (ML) DOC (mean [DOC]ML-observed = 66 /^M, mean A^^C-DOCML-observed = —230%o) is composcd of a combination of deep DOC (mean [DOCJdeep-observed = 43 /xM, mean A^'^C-DOCdeep-observed = -390%o) plus DOC assumed to be derived from new contemporary planktonic production (mean [ D O C l n e w estimated = 2 3 / x M [ = 6 6 m i u U S 4 3 / x M ] , m e a n A ^ ^ C - D O C n e w estimated

=

+116%o - the same as Sargasso Sea surface A^'^C-DIC values at the time of this study). Using the following mass-balance calculation, ([DOCIML -observed *A

C-DOCML-calculated)

= ([DOCldeep -observed

A

~r ( [ A - ' ^ ^ l n e w estimated

C-DUL-deep-observed) A

C ~ D U C n e w estimated)?

L^l

the A^'^C for mixed-layer DOC was calculated (A^'^C-DOCML-caicuiated) to be — 214%o, which is close to the average A^^C-DOCjviL-observed value of -230%o (Bauer et al, 1992a; Druffel et al, 1992). In other words, this twocomponent model, whereby old, ^"^C-depleted deep ocean DOC mixes with young ^"^C-enriched DOC from surface production, appears to describe adequately the average, relatively ^"^C-depleted values, of surface-ocean DOC. Performing the same operation for the central North Pacific (Fig. 9B) and using Eq. [4], WilHams and Druffel (1987) and Druffel et al (1992) predicted that the average mixed layer DOC (mean [DOC]ML-observed = 80 /iM, mean A^'^C-DOCML-observed = -153%o) is composcd of a combination of deep DOC ( m e a n [DOCJdeep-observed =

36 /xM, m e a n

A^'^C-DOCdeep-observed =

-525%o)

plus DOC assumed to be derived from new contemporary planktonic (mean [DOCJnewestimated = 4 4 / x M [ = 8 0 m i u U S 3 6 fjM],

m e a n A^"^C-DOCnewestimated

=

-M47%o—the same as central North Pacific surface A^'^C-DOC values at the time of this study). Again using Eq. [4], A^'^C-DOCML-caicuiated is estimated

James E. Bauer

432

MIXED LAYER newly produced DOC

total surface DOC

Assumed: DOC = ~23|LiM Ai4c = ~+116%c

Calculated: DOC = ~66|LiM Ai^C = ~-214%o

Si

DEEPLAYER

Observed: old, DOC = -'43|LiM refractory Ai^C = ~-390%o DOC

Figure 9 Two-component mixing models for evaluating the vertical distributions of A^^C-DOC in (A) the Sargasso Sea and (B) the central North Pacific. See section V.A of text for full description.

to be —155700, which is identical within analytical error (±~6%o) to the A^^C-DOCML-observed of — 153%o. We therefore conclude that in both the north Atlantic and north Pacific central gyres that the DOC in the mixed layer consists to a first approximation of deep ocean DOC that has aged and mixed vertically over time scales of at least deep ocean mixing rates (^1500 years; Stuiver et al, 1983), and of DOC derived from completely modem organic matter synthesized by plankton in the upper ocean. This exercise also demonstrates the extremely refractory nature of the old, deep ocean DOC fraction, the ultimate fate of which is not known. However, surface ocean processes such as photochemical modification (Mopper et al, 1991; Mopper and Kieber, Chapter 9) and/or

433

Carbon Isotopic Composition of DOM MIXED LAYER newly produced DOC

total surface DOC

Assumed: DOC = ~44|xM A14C = ~+147%o

Calculated: DOC = ~80^M A14C = ~-155%o

>
r Figure 9

(Continued)

bacteria degradation (Kirchman et al, 1991; Carlson and Ducklow, 1996; Carlson, Chapter 4; Cherrier, et al, 1996; 1999) of DOC may be important factors regulating the turnover of this globally significant reservoir of organic matter.

B. DISTRIBUTIONS OF 6^^C AND A ^ ^ C OF DOC IN O C E A N M A R G I N S

The information provided by A^'^C-DOC and 5^^C-D0C may also be used to evaluate the inputs (both qualitative and quantitative) of different sources of organic matter to ocean margin waters. The more highly variable distributions

200

York R. (0 psu)

MAB primary production -[]

100 0

o U

Ches. Bay (5-25 psu)

-100

MAB shelf and shallow slope

-200 -300

SS, > 1,000m

-400 -500 -600 -29

MAB deep slope

MAB nepheloid layer, >10kD

-28

-27

-26

-25

-24

-23

-22

-21

-20

5:13r

8'^Cof DOC(o/oo) b

200

0

o

Q O CJ

MAB primary production

Ches. Bay (5-25 psu)

100

1""

1

1 ^.i^^_^

/

'^^"^^^^-^^^rp.1/'*\'•••••:

-100

^ - ^ ^

-200

Grp. 2 7^^^'^'*''''10kD^,,,^

-400

1

/

-500 -600 -29

-28

-27

-26

-25

-24

-23

-22

-21

-20

S^^Cof DOG(o/oo) C

200 100 o

0 -100

o

-200

«4—

-300

O CJ 10kD __

SS >1,00'0m

-500 -600 -29

^^^-28

-27

-26

-25

-24

-23

8^3C Of DOC (0/00)

-22

-21

-20

Carbon Isotonic Composition of DOM

435

of A^'^C and 8^^C in the ocean margins (Table IV) compared to the open ocean (Table II) suggest that the origins and sources of margin are concomitantly more diverse in the margins. Several studies have recently attempted to evaluate inputs of multiple sources and ages of organic carbon to the DOC pool of the Mid-Adantic Bight region of the western North Atlantic using A^'^C and 5^^C of both total DOC and of UDOC. The major findings of these studies are summarized here. The A^'^C and 5^^C of total DOC was measured in shelf and slope waters of the Mid-Atiantic Bight by Bauer et al (2001). These data, collected between Nantucket and Cape Hatteras, show a high degree of covariance between A^'^C, (5^^C, and salinity in both shelf and shallow slope waters (Figs. 8a-8c); these relationships, however, were not found to hold for deeper slope waters (>^300 m depth; Fig. 8b), indicating at least two distinct classes of DOC. The paired A^'^C(5^^C distributions for Mid-Atlantic Bight DOC may be evaluated along with the ranges in A^'^C and (5^^C of all of the potential sources of organic matter to the DOC pool that have been measured for this region (Fig. 10a). Since the two classes of measured paired A^'^C-DOC and 5^^C-D0C values He within those of the potential sources, isotopic mass balances can be used to estimate the relative contributions from each of these sources. In simple or well-constrained systems, single isotope linear mixing models are often adequate for first-order approximations of the sources contributing to a sample. However, the number and isotopic variability of autochthonous and allochthonous sources of organic matter in margin and other coastal environments is much greater than in many other aquatic environments. In such systems, the use of multiple natural isotopes may provide a greater degree of differentiation between multiple sources than single isotopes (Williams et al, 1992). Since both A^^C and 5^^C were measured in this study, we may use a dual-isotope approach, which should provide a greater degree of specificity for organic carbon than for organic matter in general. The potential sources of DOC to MAB shelf and slope waters for which paired A^'^C and 5^^C information are available are shown in Table V and plotted in Fig. 10a (means and ranges). These sources include: (a) total freshwater DOC from Chesapeake Bay, as represented by one of its major subestuaries, the

Figure 10 (a) Mean values and ranges in A^^C and 6^^C of potential sources of DOC to MidAtlantic Bight (MAB) shelf and slope waters. Isotope values for potential sources were obtained from the following: York River total DOC—Raymond and Bauer (2001b) and Raymond and Bauer (2001c); Chesapeake Bay > 1-kDa material and MAB > 10-kDa nepheloid layer material—Guo et al. (1996); Sargasso Sea (SS) total DOC—Bauer et al (1992a) and Druffel et al (1992); MAB primary production—based on A^^C and ^^^C values for DIC in Bauer et al (2001). (b) The A^^^C vs ^^^C fields of potential source combinations of DOC to MAB shelf and shallow slope waters only and (c) MAB deep slope waters only. See section V.B. of text for details. Adapted with permission from Bauer e/fl/. (2001).

436

James E. Bauer Table V

Mean A^^C and S^^C Values of Potential DOC Sources Used to Calculate Relative Contributions to DOC in the Mid-Atlantic Bight (Refer to Fig. 10). Potential Source York River freshwater DOC

Mean A^^C (%o)

Mean ^^^C i%o)

200

-28.4

Raymond and Bauer, 2001c

Reference

Chesapeake Bay (5-25 psu) HMW(>lkDa)DOC

-1

-27.8

Guo etal, 1996

Deep MAB VHMW (> 10 kDa) DOC

-580

-26.4

Guo etal, 1996

55

-20^

Bamretal,

-394

-20.8

BmcTetaL, 1992a

-238

-21.2

Dmffe\ et ai, 1992

MAB primary production'' Deep Sargasso DOC Surface Sargasso DOC

2001

Note. Adapted from Bauer et al. (2001). ^ Based on A^^C and ^^^C of DIC (Bauer et al, 2001). ^Assumes fractionation of -19%o during CO2 fixation by marine phytoplankton.

York River estuary (Raymond and Bauer, 2001c), (b) the high molecular weight (> 1 kDa) component of DOC from the mainstem of Chesapeake Bay (Guo et al, 1996), (c) the very high-molecular-weight (VHMW, > 10 kDa) component of MAB near-bottom nepheloid material (Guo et al, 1996), (d) present-day primary production in MAB surface waters, estimated from the A^'^C and 5^^C values of DIC, and (e) previous estimates of fully marine values for the surface and deep Sargasso Sea (Bauer et al, 1992a; Druffel et al, 1992), taken to be representative of open North Atlantic waters in general. It is possible that the two different molecular weight fractions (>1 kDa and >10 kDa) used in this analysis vary isotopically from the bulk DOC from the same ocean margin waters. However, without further comparative isotopic information between the high-molecular-weight components and bulk DOC, we must assume for the present exercise that they are comparable in isotopic content, similar to the observations of Bauer et al (submitted for publication) for open ocean waters (see Table II). Use of the isotopic values of Chesapeake Bay >1 kDa and deep MAB > 10 kDa fractions of DOC (Guo et al, 1996) as terrestrial/riverine end-members is supported by thefindingsof Mitra^fa/. (2000) who showed that both of these components contained elevated amounts of lignin-derived phenols originating from terrestrial plants. The paired A^'^C and 5^-^C distributions of DOC samples measured in this study (Fig. 10a) are consistent with DOC in the MAB being composed of one or more of the indicated potential sources. In addition, A^'^C and 5^^C distributions for DOC from MAB shelf and shallow slope waters lie between different potential

437

Carbon Isotopic Composition of DOM

end-members (i.e., surface Sargasso, Chesapeake BayAfork River and MAB primary production) than does DOC from deeper slope waters (i.e., deep Sargasso and MAB nepheloid layer material). We hypothesize that this is a result of unique sources (and ages) of DOC to these two major water types. It is also possible from Fig. 10a for admixtures of MAB modem primary production and > 10-kDa nepheloid material to give A^'^C and 8^^C values similar to those observed in MAB shelf and shallow slope waters. However, two factors argue against this proposed scenario. First, the > 10-kDa ^"^C-depleted material in deeper waters comprises only 3-6% of the total DOC (Guo et al, 1996) and second, the observations of Guo et al. (1996) indicate that in shallow waters of the MAB, the > 10-kDa fraction was actually similar in both A^'^C and 8^^C to values for MAB primary production (Fig. 10a). Furthermore, as demonstrated by Druffel et al. (1992), Sargasso Sea surface ocean DOC can be adequately described as a combination of old, deep material and young material from primary production. For purposes of the following exercise, we assume that the major inputs to the different shelf and surface slope waters (shown as Groups 1, 2, and 3 in Fig. 10b) can be described reasonably by a combination of (a) surface Sargasso Sea DOC (which itself is composed of deep Sargasso material and recent marine primary production; Bauer et al, 1992a; Druffel et al, 1992), (b) Chesapeake Bay DOC (which must also contain some York River material), and (c) DOC derived from contemporary MAB primary production. In order to establish first-order estimates of the relative contributions of the major presumed sources of DOC to shelf and shallow slope waters, we used three-source isotopic mixing models similar to those of Fry and Sherr (1984) and Kwak and Zedler (1997). The generaUzed mixing equation is ^DOC-MAB = / l ^ D O C - / l + fl^DOC-f2

+ (1 " / l " /2)^DOC-/3.

[5]

where X is the isotopic composition (A^'^C and 8^^C) of DOC from the MAB observed. The value / is the relative contribution of each of the three potential sources to the total DOC in the MAB samples, and/i-|-/2 +/3 = 1.0. Since there are two unknowns (/l and/2) in Eq. (1), the equation must be solved simultaneously using A^'^C and 8^^C. The contribution of the third potential source,/s, is equal to(l-/i-/2).

The results of these calculations (Table VI) indicate that shelf and shallow slope waters are dominated (up to 97%) by DOC that is similar isotopically to that found in the open North Atlantic (Sargasso Sea). However, the DOC from different regions within the MAB contains varying and often significant amounts of material from in situ production and material that must arise from terrestrial, riverine, and/or estuarine (TRE) inputs. For Group 1 (Fig. 10b; Table VI), up to a third of the DOC is TRE material, and sHghtly more (25^3%) is recently derived MAB primary production. Group 2 samples (Fig. 10b) are composed of lower amounts of DOC derived from both TRE (9-25%) as well as MAB primary production (0-12%) (Table VI). Finally, the two anomalous samples (Group 3) that

438

James E. Bauer

Table VI Estimates of Relative Inputs of Different Potential Sources of DOC to the Middle Atlantic Bight, Based on Results of Three-Source Isotopic Mixing Models Relative contribution (%) of: Zone Shelf and shallow slope Group l'^ Group 2" Group 3" Deep slope^

Ches./ York

MAB prim, prod.

Sargasso shallow

Sargasso deep

MAB nepheloid

MAB surf, sediments

19-31 9-25 2-3

25-43 0-12 na

26-52 64-97 77-85

na na na

na na 12-21

na na na

0-3

na

na

74-88

8-25

na

Note. See text for details. Adapted from Bauer et al. (2001). na, not applicable (end-member not used in mass balance). ^ As shown in Fig. 10b. ^ As shown in Fig, 10c. lie outside of the mixing fields that contribute to the majority of shelf and shallow slope samples (Fig. 10b) must have a component that is much older in order to account for the observed values. The only material that has been identified that can fulfill the requirement of a simultaneously ^"^C- and ^-^C-depleted DOC component is the very high-molecular-weight DOC (>10 kDa) from the nepheloid layer (Guo et al, 1996; Guo and Santschi, 2000). When this source is mass-balanced against shallow Sargasso and TRE material, we find that it comprises 12-21% of the total DOC (Table VI), while younger TRE material represents only trace (2-3%) inputs. Following our approach for shelf and shallow slope waters, we assume that MAB deep slope DOC is composed predominantly of a combination of deep Sargasso, >10-kDa nepheloid, and TRE material (Fig. 10c). Similar to the Group 3 DOC samples (Fig. 10b), we find that up to 25% of the deep slope DOC may be composed of a presumably highly aged, ^^C-depleted high-molecular-weight component (Table VI). On the basis of several other recent studies (Druffel and Williams, 1990;Sherrell^M/., 1998;Bianchiert2/., 1998; Bauer and Druffel, 1998; Druffel et al, 1998; Honda et al, 2000), lateral inputs of organic matter from both the nepheloid layer and sediments in continental margins are plausible sources of organic matter not only to slope waters, but to even more oceanic waters. The A ^"^C-POC values in MAB slope waters are even more highly depleted compared to the open North Atlantic (Bauer etal, 2001), suggesting that margins are a source of ^"^C-depleted POC (and by similar reasoning, DOC) to the open ocean water column. The fact that the high-molecular-weight component has substantial terrestrial 8^^C character and is concomitantly so old (Guo et al, 1996; Guo and Santschi, 2000), suggests that slope sediments could represent a temporary "aging

Carbon Isotopic Composition of DOM

439

reservoir" for terrestrial and shelf/slope-derived organic matter. This material, possibly deposited initially in slope and certain shelf sediments as POC or mineralsorbed DOC (Mayer, 1994; Keil et al, 1997), may then undergo partial postdepositional desorption, hydrolysis, and degradation in sediments prior to being re-released to the water column pool of DOC (Burdige and Gardner, 1998; Burdige et al, 1999; Alperin et a/., 1999; Burdige, Chapter 13). If it occurs, this proposed mechanism is significant in that it would result in margin sediments providing a source of "pre-aged" terrestrial and shelf/slope primary production to the deep ocean directly (Bauer et al, 1995). Finally, on the basis of both past and recent evidence (Spiker and Rubin, 1975; Raymond and Bauer, 2001b), we cannot, without further information, rule out the possibility that rivers themselves transport directly a significant amount of aged terrestrial DOC to certain coastal systems. For example, Spiker and Rubin (1975) reported A^'^C values for total DOC in the Rappahannock and Susquehanna Rivers of —91 and —81%o, respectively, while Raymond and Bauer (2001b) have found mean A^'^C-DOC values of -158%o in the freshwater Hudson River.

C. SOURCES AND INPUTS OF U D O C TO OCEAN MARGINS In addition to being used to track sources and inputs of bulk DOC, natural ^"^C and ^^C have also been applied for tracing the origins and ages of UDOC in the Mid-Atlantic Bight and Gulf of Mexico shelf and slope regions. The A^'^C and 8^^C signatures of various molecular weight fractions of UDOC (primarily the > 1 -kDa and > 10-kDa fractions) appear to be more variable than, and differentiable from, bulk DOC from the same environment, and between different environments. Similar to total DOC, Santschi et al. (1995) and Guo et al. (1996) observed inverse correlations between A ^ ^ ^ . U ^ Q C and ^^^C-UDOC in the Gulf of Mexico and Mid-Atlantic Bight margins (Figs. IIA and UB); both A^^C-UDOC and 5^^C-UDOC also correlated inversely with salinity, suggesting a possible application of these relationships for identifying sources of different aged terrestrial and marine (including estuarine and sedimentary UDOM and shelf/slope production) to the UDOC, and hence total DOC, pools. These workers also employed the relationship between the C:N ratios within UDOC (or more appropriately, UDOM, or ultrafiltered dissolved organic matter) to distinguish different end-member "classes" of UDOM and use them to estimate their contributions to the observed A^'^C and C:N values in these two regions (Figs, l i e and IID). In the Gulf of Mexico, UDOM was observed to consist primarily of one of three end-members (deep-water, offshore surface or estuarine; Fig. IIC), with little mixing between the three. In contrast, Mid-Atlantic Bight UDOM appeared to consist to a much larger extent of admixtures between either (a) deep-water and offshore surface UDOM or (b) offshore surface and estuarine

440

James E. Bauer

H 1A +

1

o

Deepwatei colloids

T +

/ /

T

1

— h -

—1

1

Estuarine colloids

o

o

I oo\ 1 /o 1 V/o / '

0

N^V

\——1

"T

-]-

.o y 1 —\—

-200 * C

4-4-

^^ooV

Offshore surface\ water colloids

-1

h

J_

1-

\

-100

(%«)

(%o)

150 L • • • 1 • • ' i • • • 1 • • • 1 • • • 1 • • • J

H

J_B -|-

\

h -—f

H—H

1

b

1

100

i ^

Middle Atlantic Bight (Surface water COM^)

E

50

/ ^

Deep water colloids

T ^m^

4- V-X""**^*

-f -1

1

Estuarine colloids

• •^ 1

-500

h-—1

-400

-300

A" C

\

/ A

o

1 T

" • J Offshore surface 1 -^ water colloids T

-H—

H -200 -100 (%o)

1

1

h

0 -50

1 ^^^"^O i

T

-150

t

-200

f

-100

t

A-^A

]

l

j T

A

A^ i ^ ^^^ ^^"^--^ ^^ I / t

(R=0.71, n=14)

A

+

-250 r • • ' 1 • • • 1 • • • 1 • ' ' 1 ' • ' 1 ' • • 1 -32 -28 -26

5 " C (%c)

Figure 11 Relationships between: (A) C/N and A^'^C of UDOC for the Gulf of Mexico, (B) C/N and A^^^C of UDOC for the Mid-Atlantic Bight, (C) A^^C-UDOC and ^^^C-UDOC for the Gulf of Mexico, and (D) A^^C-UDOC and ^^^C-UDOC for the Mid-Atlantic Bight. COMi refers to >1 kDa UDOM: Adapted with permission from Guo et al. (1996).

UDOM (Fig. 1 ID), but not between deep-water and estuarine UDOM. Thus, lowsalinity sources of both UDOC and total DOC (see Section IV.B, above) appear to be isotopically discernible for a considerable distance from riverine and estuarine sources in ocean margin regions. This further emphasizes that different ocean margin regions may be unique from one another with respect to their sources and inputs of organic matter, depending upon the relative magnitudes of terrestrial and marine fluxes, and with respect to the hydrographic (i.e., mixing) features of a particular ocean margin region. These same workers (Guo et al., 1996) used the A^'^C of > 1 kDa and > 10 kDa and total DOC (based on independent measurements by Bauer et al., 2001) to estimate by mass-balance the relative contributions and A^^C signatures of < 1-kDa, 1- to 10-kDa, and > 10 kDa material to the standing stock of total DOC in the Mid-Atlantic Bight shelf and slope off Cape Hatteras. Results of these mass balance calculations (Table VII) indicate that very low-molecular-weight UDOC ( 10-kDa fractions. Furthermore, the < 1-kDa fraction

441

Carbon Isotopic Composition of DOM Table VII Estimates of Relative Inputs and A^'^C Signatures of Different Molecular Weight Fractions of UDOC to the Middle Atlantic Bight, Based on Results of a Mass Balance Mixing Model Relative contribution of

Stn Shelf 10 Slope 1 12 13

Depth (m)

10kDa

1-10 kDa

%^

Al^C^

%

Ai^C^

%

2 25

-257 -487

66 69

-128 -334

23 27

-6 -132

11 4

2 750 2 2300 2 250 2600

-217 -442 -240 -451 -246 -443 -452

65 70 66 71 66 71 72

-175 -399 -143 -359 -132 -355 -336

30 26 29 26 28 24 25

-160 -427 -9 -558 -8 -611 -709

5 4 5 3 6 5 3

Note. See text for details. Adapted from Guo et al. (1996). ^ C a l c u l a t e d a s A ^ ^ ^ ^ ^ ^ ^ = (Al'^CtotalDOC - ([Al'^Cl-lOkDa X %l-lOkDa] + [A^'^C>iOkDa X Al4C>i0kDa])/% 1-kDa fraction and the total DOC in open ocean settings (see Table II, for eastern North Pacific UDOC and total DOC). Thus, size-fractionation of DOC and isotopic analysis of UDOC fractions provides an additional tool for assessing sources and contributions of DOC in ocean margin regions. D. EXCHANGES OF DOC BETWEEN THE OCEAN'S MARGINS AND

ITS INTERIOR

As illustrated in Fig. 7, gradients exist between the A^'^C-DOC signatures of certain ocean margins (e.g., western North Atlantic and eastern North Pacific) and the contiguous open ocean. It has also been shown by Bauer and Druffel (1998) that both of these margins have several micromolar higher average deep DOC concentrations than more remote ocean regions such as the Sargasso Sea and central

442

James E. Bauer

North Pacific. The origin(s) of this old, ^"^C -depleted carbon to continental slope and rise waters is (are) not known, but several possibilities may be considered. In the western North Atlantic, the reintroduction to the water column of old sedimentary organic carbon (initially as both DOC and POC) from weathered shelf and upper slope sediments (Anderson et ai, 1994; Churchill et al, 1994), porewaters (Bauer et ai, 1995; Burdige, Chapter 13), and even submarine hydrocarbon seeps (Boehm and Requejo, 1986; Roberts and Carney, 1997) may contribute to the highly ^"^C-depleted (A^'^C as low as ^—700%o) colloidal and dissolved organic carbon observed in near-bottom waters in the Middle Atlantic Bight (Guo et ai, 1996) as well as to the ^"^C-depleted suspended POC observed in the water column there (Bauer et ai, 2001). The elevated DOC concentrations in slope (in the westem North Atlantic) and rise (in the eastern North Pacific) waters, along with lower A ^"^C-DOC values, indicate that DOC is present in ocean margins that is both older than that in the North Atlantic and Pacific central gyres and potentially available for export to the open ocean (Wollast, 1991, 1998). As a result of these gradients, simple two-source box models have been employed to assess the relative contributions to the deep, interior ocean DOC pools from (a) margin-derived DOC and (b) surface ocean-derived DOC from contemporary production (Bauer and Druffel, 1998). The conceptual framework and the relevant mass-balance relationships used for this assessment are shown in Fig. 12. The presence of positive concentration gradients between the margins and deep open ocean and between the surface and deep open oceans indicate that both

inputs from primary & secondary production

I

A^'^C: greater than deep ocean [DOC]: greater than deep ocean A^'^C: less than deep ocean [DOC]: greater than deep ocean

assumed steady-state deep ocean MASS BALANCE: (A-A i^C-DCX: X A[DOC])^^„,^^^, + (A-A •^C-DOC x A[DOC]),^^,„„p^en, = (A-A i^C-DOC X A[DOC])„id.gy^ = 0 at steady-state

Figure 12 Conceptual model of two-component steady-state inputs of margin-derived and surfaceocean-derived DOC to the deep open ocean. See section V.D of text for details.

Carbon Isotonic Composition of DOM

443

margins and the surface ocean may represent sources of DOC to the deep central gyres (Table VIII). Bauer and Druffel (1998) estimated by ^"^C mass balance the relative potential contributions of each of these sources to the deep North Atlantic and Pacific using the following simplifying assumptions: (a) the deep central North Atlantic and Pacific are in steady state with respect to A^'^C values and concentrations of DOC (Williams and Druffel, 1987; Bauer et al, 1992a; Druffel et al, 1992); (b) the two dominant sources of DOC to the deep central gyres are lateral inputs of ^^C-depleted material derived from the margins and vertical inputs of "modem," ^^C-enriched material derived from surface ocean production (Druffel et al, 1996); and (c) the margin-to-deep open ocean and surface-to-deep open ocean gradients observed in these studies are representative of the North Atlantic and Pacific as a whole. We find that in order to maintain the observed average A^^C-DOC values in the deep central gyres, the input of DOC from the margins is calculated to be as much as 25-100 times that of modem, surface ocean-derived carbon (Table VIII). These estimates of margin and surface ocean contributions to the deep open ocean have two main implications. First, inputs of "aged" DOC from the margins to the deep open ocean may surpass inputs derived from recent surface ocean production. Second, in view of the much larger surfaceto-deep vs margin-to-deep concentration gradients, the vast majority of young, surface-derived material must be degraded, allowing a smaller but more highly refractory margin component to contribute proportionally more to the deep central gyres. Although no a priori assumptions are made in these estimates about the specific mechanisms promoting horizontal exchanges from the margins, transport of ^^C-depleted DOC from ocean margins to the central gyres may be facilitated by isopycnal (i.e., lateral) eddy diffusion, which can be 10^-10^ times greater than vertical eddy diffusive transport (Knauss, 1978). The isotopic signatures of DOC at the coastal/open ocean boundaries (i.e., slope and rise waters) indicate that this carbon has a mainly nonrecent marine origin and is older than organic carbon from the North Atlantic and North Pacific central gyres. If this material propagates seaward, possibly along isopycnal surfaces, it may represent a source of old DOC to intermediate and deep waters of the interior ocean (WoUast, 1991,1998). Alternative mechanisms such as overtuming circulation in regions of intermediate and deep-water formation have been proposed for transporting surface ocean DOC to the deep central gyres and are discussed in detail in Hansell and Carlson (1998), Hansen et al in press) and Hansell (Chapter 15).

VL SUMMARY AND FUTURE CHALLENGES This review and synthesis of the available information on the isotopic (^"^C and ^^C) composition of DOC in open-ocean and ocean-margin environments demonstrates that both of these isotopes can provide useful information on the sources

I e

o

t o U

I

'& -2

5 2

til

-g

II

8

I

O

d

P:

I

>n

d

S S

< o

Carbon Isotopic Composition of DOM

445

and ages of DOC. ^"^C and ^^C constitute powerful tracers of DOC sources and cycling times in both types of environment, especially when used in conjunction with one another, or with other source and age parameters (e.g., salinity, C/N, A^^^C-DIC, etc.). In spite of the invaluable utility of ^"^C and ^^C in studies of ocean DOC cycling, the vast majority of measurements to date have been made on bulk or total pools of DOC, or on operationally defined subfractions such as humic substances or UDOM. There is presently a fundamental need for information in marine organic isotope geochemistry on the sources and ages of DOC to oceanic and coastal waters, and for understanding the internal factors responsible for altering the isotopic and biochemical composition of DOC in marine waters during its transformation and aging. In order to obtain more information on these topics, we propose the following suggested areas of future research on the natural isotopic characterization of DOC in the oceans: (i) Assessing ^"^C and ^^C in autochthonous marine sources of DOC, including, but not limited to, living planktonic biomass, sediment porewaters, and solubilized and degrading sinking POC (e.g., from sediments trap studies). (ii) Evaluating the magnitudes of allochthonous inputs on oceanic distributions of ^"^C and ^^C, including, but not limited to, terrestrial inputs and atmospheric deposition (including the role of black carbon and both natural and anthropogenic hydrocarbons); better characterization of terrestrial, riverine, and estuarine isotopic source signatures; and alterations in the ^"^C and ^^C signatures of terrestrial and riverine DOC in estuaries and the coastal ocean. (iii) Studies of the effects of both biotic and abiotic factors controlling ^"^C and ^^C distributions in DOC. Such factors include changes in ^"^C and ^^C contents during DOC degradation by heterotrophic bacteria due to both preferential utilization of more labile constituents and isotopic fractionation and the potential role of abiotic factors such as sorption, desorption, photolysis, etc. (iv) Evaluating inputs of young, labile DOC to the deep ocean by "shortcircuiting" due to intermediate and deep water formation. (v) Further partitioning and isotopic characterization of the constituents of DOC and UDOC, utilizing compound-class and compound-specific separation techniques and ^^C and ^^C isotopic analyses. Although these techniques have so far been applied successfully to studies of sedimentary organic carbon and POC, they have only recently been extended to studies of marine DOC cycling (Aluwihare, 1999), and it is anticipated that there will be an expansion of such studies in the near future.

ACKNOWLEDGEMENTS I thank numerous individuals for their encouragement, hard work, and coUegiality during the course of this research over the years, foremost among them Drs. Peter M. WilHams and Ellen

446

James E. Bauer

R. M. Druffel. Others in our research groups who made possible our own work in the area of organic isotope geochemistry include Dave Wolgast, Ken Robertson, Sheila Griffin, Pete Raymond, Ai Ning Loh, Jennifer Cherrier, Carrie Masiello, Mark Schrope, and Eva Bailey. Michaele Kashgarian and John Southon of the Center for Accelerator Mass Spectrometry, Lawrence Livermore National Laboratory, were instrumental in making available resources and facilities for AMS A^^C analyses, and Eben Franks performed S^^C analyses. I also thank the captains and crews of a number of UNOLS vessels for helping with the logistics necessary to conduct the fieldwork, including RA^'s Melville, Knorr, New Horizon, Seward Johnson, Endeavor, Columbus Iselin, and others. This work was supported primarily by the Chemical Oceanography Program of the U.S. National Science Foundation and the U.S. Department of Energy's Ocean Margins Program.

REFERENCES Aiken, G. R., Thurman, E. M., Malcolm, R. L., and Waltom, H. F (1979). Comparison of XAD macroporous resins for the concentration of fulvic acids from aqueous solution. Anal. Chem. 51, 17991803. Alperin, M. J., Martens, C.S., Albert, D. B., Suayah, L B., Benninger, L. K., Blair, N. K., and Jahnke, R. A. (1999). Benthic fluxes and porewater concentration profiles of dissolved organic carbon in sediments from the North Carolina continental slope. Geochim, Cosmochim. Acta 63,427^88. Aluwihare, L. L (1999). "High Molecular Weight Dissolved Organic Matter in Seawater: Chemical Structure, Sources and Cychng." Ph.D. thesis, Massachusetts Institute of TechnologyAVoods Hole Oceanographic Institution Joint Program in Oceanography, Cambridge, MA. Anderson, R., Rowe, G., Kemp, P., Trumbore, S., and Biscaye, P. (1994). Carbon budget for the mid-slope depocenter of the Middle Atlantic Bight. Deep-Sea Res. 41,669-703. Arevena, R., and Wassenaar, L. I. (1993). Dissolved organic carbon and methane in a regional confined aquifer, southern Ontario, Canada: Carbon isotope evidence for associated subsurface sources. Appl Geochem. 8,483-493. Armstrong, F. A. J., WiUiams, P. M., and Strickland, J. D. H. (1966). Photo-oxidation of organic matter in seawater by ultra-violet radiation, analytical and other apphcations. Nature 211,481^83. Bacastow, R. B., Keeling, C. D., Lueker, T J., Wahlen, M., and Mook, W. G. (1996). The S^^C Suess effect in the world surface oceans and its implications for oceanic uptake of CO2. Global Biogeochem. Cycles 10, 335-346. Bauer, J. E., and Druffel, E. R. M. (1998). Ocean margins as a significant source of organic matter to the deep ocean. Nature 392,482-485. Bauer, J. E., Druffel, E. R. M., Williams, PM., Wolgast, D. M., and Griffin, S. (1998a). Temporal variability in dissolved organic carbon and radiocarbon in the eastern North Pacific Ocean. J. Geophys. Res. 103(C2), 2,867-2,882. Bauer, J. E., Druffel, E. R. M., Wolgast, D. M., and Griffin, S. (2001). Cycling of dissolved and particulate organic radiocarbon in the northwest Atlantic continental margin. Global Biogeochem. Cycles 15(3): 615-636. Bauer, J. E., Haddad, R., and DesMarais, D. (1991). Method for the determination of stable carbon isotope ratios in dissolved organic matter of interstitial and other natural marine waters. Mar Chem. 33,335-351. Bauer, J. E., Reimers, C. E., Druffel, E. R. M., and Williams, P. M. (1995). Isotopic constraints on carbon exchange between deep ocean sediments and seawater. Nature 373,686-688. Bauer, J. E., Ruttenberg, K. C , Wolgast, D. M., Monaghan, E., and Schrope, M. K. (1996). Crossflow filtration of dissolved and colloidal nitrogen and phosphorus in seawater: results from an intercomparison study. Mar Chem. 55, 33-52.

Carbon Isotonic Composition of DOM

447

Bauer, J. E., Williams, P. M., and Druffel, E. R. M. (1992a). ^"^C activity of dissolved organic carbon fractions in north-central Pacific and Sargasso Sea. Nature 357, 667-670. Bauer, J. E., Williams, P. M., and Druffel, E. R. M. (1992b). Recovery of sub-milligram quantities of carbon dioxide from gas streams by molecular sieve for subsequent determination of isotopic (^^C and ^^C) natural abundances. Analyt. Chem. 64, 824-827. Bauer, J. E., Druffel, E. R. M., Wolgast, D. W, and Griffin, S. (submitted for publication). Radiocarbon in colloidal and subcoUoidal organic matter in the open ocean. Geophys. Res. Lett. Bauer, J. E., Wolgast, D. W, Druffel, E. R. M., Griffin, S., and Masiello, C. A. (1998b). Distributions of dissolved organic and inorganic carbon and radiocarbon in the eastern North Pacific continental margin. Deep-Sea Res II45, 689-714. Benner, R. (1991). Ultrafiltration for the concentration of bacteria, viruses, and dissolved organic matter. In "Marine Particles: Analysis and Characterization" (D.C. Hurd and D.W. Spencer, Eds.), Geophysical Monograph 63, pp. 181-185. American Geophysical Union, Washington, DC. Benner, R. (2002). Chemical composition and reactivity. In "Biogeochemistry of Marine Dissolved Organic Matter" (D.A. Hansell and C.A. Carlson, Eds.), pp. 59-90. Academic Press, San Diego. Benner, R., Biddanda, B., Black, B., and McCarthy, M. (1997). Abundance, size distribution, and stable carbon and nitrogen isotopic composition of marine organic matter isolated by tangentialflow ultrafiltration. Mar. Chem. 57, 243-263. Benner, R., Chin-Leo, C , Gardner, W, Eadie, B., and Cotner, J. (1992b). "The Fates and Effects of Riverine and Shelf-Derived DOM on Mississippi River Plume/Gulf Shelf Processes, Nutrient Enhanced Coastal Ocean Productivity, NECOP Workshop Proceedings, pp. 84-94. Texas A& M University College Sea Grant Program, College Station, TX. Benner, R., Pakulski, J. D., McCarthy, M., Hedges, J. I., and Hatcher, R G. (1992a). Bulk chemical characteristics of dissolved organic matter in the ocean. Science 255,1,561-1,564. Bianchi, T. S., Bauer, J. E., Druffel, E. R. M., and Lambert, C. D. (1998). Pyrophaeophorbide-a as a tracer of suspended particulate organic matter from the NE Pacific continental margin. Deep-Sea Res. II 45(4-5), 115-131. Bianchi, T. S., Lambert, C. D., Santschi, P. H., and Guo, L. (1997). Sources and transport of landderived particulate and dissolved organic matter in the Gulf of Mexico (Texas shelf/slope): The use of lignin-phenols and loliolides as biomarkers. Org. Geochem. 27, 65-78. Boehm, P. D., and Requejo, A. G. (1986). Overview of the recent sediment hydrocarbon geochemistry of Atlantic and Gulf coast outer continental shelf environments. Estuar Coastal Shelf Sci. 23, 29-58. Boutton, T. W (1991a). Stable carbon isotope ratios of natural materials. L Sample preparation and mass spectrometric analysis. In "Carbon Isotope Techniques" (D.C. Coleman and B. Fry, Eds.), pp. 155-171. Academic Press, New York. Boutton, T. W (1991b). Stable carbon isotope ratios of natural materials IL Atmospheric, terrestrial, marine and freshwater environments. In "Carbon Isotope Techniques" (D.C.Coleman and B. Fry, Eds.), pp. 173-185. Academic Press, New York. Broecker, W. S. (1991). The great ocean conveyor. Oceanography 4(2), 79-89. Broecker, W S., Gerard, R., Ewing, M., and Heezen, B. C. (1960). Natural radiocarbon in the Atlantic Ocean. /. Geophys. Res. 65, 2903-2931. Broecker, W S., and Peng, T.-H. (1982). "Tracers in the Sea." LDGEO Press, New York. 690 pp. Broecker, W. S., Peng, T.-H., Ostlund, G., and Stuiver, M. (1985). The distribution of bomb radiocarbon in the ocean. /. Geophys. Res. 90(C4), 6953-6970. Broecker, W S., Sutherland, S., Smethie, W, Tsung-Hung, P., and Ostlund, G. (1996). Oceanic radiocarbon: Separation of the natural and bomb components. Oceanogr. Lit. Rev. 43,28. Brophy, J. E., and Carlson, D. J. (1989). Production of biologically refractory dissolved organic carbon by natural seawater microbial populations. Deep-Sea Res. 36,497-507.

448

James E. Bauer

Buesseler, K. O., Bauer, J. E., Chen, R. R, Eglinton, T. I., Gustafsson, O., Landing, W., Mopper, K., Moran, S. B., Santschi, P. H., Vemon-Clark, R., and Wells, M. L. (1996). An intercomparison of cross-flow filtration techniques used for sampling marine colloids: Overview and organic carbon results. Mar. Chem. 55, 1-32. Burdige, D. J. (2002). Sediment pore waters. In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. Hansell and C. A. Carlson, Eds.), pp. 611-663. Academic Press, San Diego. Burdige, D. J., Burrelson, W. M., Coale, K. H., McManus, J., and Johnson, K. S. (1999). Huxes of dissolved organic carbon from California continental margin sediments. Geochim. Cosmochim. Acto 63, 1507-1515. Burdige, D. J., and Gardner, K. G. (1998). Molecular weight distribution of dissolved organic carbon in marine sediment pore waters. Mar. Chem. 62,45-64. Carlson, C. A. (2002). Production and removal processes. In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. Hansell and C. A. Carlson, Eds.), pp. 91-151. Academic Press, San Diego. Carlson, C. A., and Ducklow, H. W. (1996). Growth of bacterioplankton and consumption of dissolved organic carbon in the oligotrophic Sargasso Sea. Aquat. Microb. Ecol. 10,69-85. Carlson, C. A., Ducklow, H. W., and Michaels, A. F. (1994). Annual flux of dissolved organic carbon from the euphotic zone in the northwestern Sargasso Sea. Nature 371,405^08. Cherrier, J., Bauer, J. E., and Druffel, E. R. M. (1996). Utihzation and turnover of labile dissolved organic matter by bacterial heterotrophs in eastern North Pacific surface waters. Mar. Ecol. Prog. Ser. 139,267-279. Cherrier, J., Bauer, J. E., Druffel, E. R. M., Coffin, R. B., and Chanton, J. C. (1999). Radiocarbon in marine bacteria: Evidence for the ages of assimilated carbon. Limnol. Oceanogr. 44,730-736. Churchill, J., Wirick, C , Flagg, C , andPietrafesa, L. (1994). Sediment resuspension over the continental shelf east of the Delmarva Peninsula. Deep-Sea Res. 41, 341-364. Clercq, M. le, van der Plicht, J., and Meijer, H. A. J. (1998). A supercritical oxidation system for the determination of carbon isotope ratios in marine dissolved organic carbon. Anal. Chim. Acta 370, 19-27. Cole, J. J., Likens, G. E., and Strayer, D. L. (1982). Photosynthetically produced dissolved organic carbon: An important carbon source for planktonic bacteria. Limnol. Oceanogr 27,1080-1090. Craig, H. (1953). The geochemistry of the stable carbon isotopes. Geochim. Cosmochim. Acta 3,53-92. Craig, H. (1971a). The deep metabolism: O2. J. Geophys. Res. 76, 299-316. Craig, H. (1971b). Son of abyssal carbon. J. Geophys. Res. 76(21), 5,133-5,139. Degens, E. T, Guillard, R. R. L., Sackett, W. M., and Hellebust, J. A. (1968). Metabolic fractionation of carbon isotopes in marine plankton. L Temperamre and respiration experiments. Deep-Sea Res. 15,1-9. Druffel, E. R., and Williams, P. M. (1990). Identification of deep marine source of particulate organic carbon using bomb ^"^C. Nature 347, 172-174. Druffel, E. R. M., and Bauer, J. E. (2000). Radiocarbon distributions in Southern Ocean dissolved and particulate organic matter. Geophys. Res. Lett. 47, 1495-1498. Druffel, E. R. M., Bauer, J. E., Williams, PM., Griffin, S., and Wolgast, D. M. (1996). Seasonal variability of radiocarbon in particulate organic carbon in the northeast Pacific. J. Geophys. Res. 97,15,639-15,659. Druffel, E. R. M., Griffin, S., Bauer, J. E., Wolgast, D. M., and Wang, X.-C. (1998). Distribution of particulate organic carbon and radiocarbon in the water column from the upper slope to the abyssal northeastern Pacific Ocean. Deep-Sea Res. II 45(4-5), 667-687. Druffel, E. R. M., Williams, P. M., Bauer, J. E., and Ertel, J. (1992). Cycling of dissolved and particulate organic matter in the open ocean. J. Geophys. Res. 97,15,639-15,659. Druffel, E. R. M., Williams, P. M., and Suzuki, Y. (1989). Concentrations and radiocarbon signatures of dissolved organic matter in the Pacific Ocean. Geophys. Res. Lett., 16, 991-994.

Carbon Isotopic Composition of DOM

449

Eadie, B. J., Jeffrey, L. M., and Sackett, W. M. (1978). Some observations on the stable carbon isotope composition of dissolved and particulate organic carbon in the marine environment. Geochim. Cosmochim. Acta 42,1,265-1,269. Eglinton, T. L, Aluwihare L. I., Bauer J. E., Druffel E. R. M., and McNichol A. P. (1996). Gas chromatographic isolation of individual compounds from complex matrices for radiocarbon dating. Anal Chem. 68,904-912. Eghnton, T. E., Benitez-Nelson, B., McNichol, A., Bauer, J. E., and Druffel, E. R. M. (1997). Variabihty in radiocarbon ages of individual organic compounds from marine sediments. Science 277, 796799. Elmore, D., and Phillips, F. M. (1987). Accelerator mass spectrometry for measuremen of long-lived radioisotopes. Science 236,543-550. Ertel, J. R., Hedges, J. I., Devol, A. H., Richey, J. E., and Ribeiro, Ribeiro, M. de N. G. (1986). Dissolved humic substances of the Amazon River system. LimnoL Oceanog. 31(4), 739-754. Falkowski, P. G. (1991). Species variability in the fractionation of ^^C and ^^C marine phytoplankton. /. Plankton Res. 13 (Suppl.), 21-28. Fogg, G. E. (1983). The ecological significance of algal extracellular products of phytoplankton photosynthesis. Bot. Mar. 26, 3-14. Fritz, P., and Fontes, J. C. (1980). Introduction. In "Handbook of Environmental Isotope Geochemistry" (P. Fritz and J. C. Fontes, Eds.), pp. 1-19. Elsevier Scientific, Amsterdam. Fry, B., HuUar, S. S., and Peterson, B. J. (1993). Platinum-catalyzed combustion of DOC in sealed tubes for stable isotopic analysis. Mar. Chem. 41,187-193. Fry, B., Peltzer, E. T., Hopkinson, C. S., Jr., Nolin, A., and Redmond, L. (1996). Analysis of marine DOC using a dry combustion method. Mar. Chem. 54, 191-201. Fry, B., and Sherr, E. B. (1984). 8^^C measurements as indicators of carbon flow in marine and freshwater ecosystems. Cort^nZ?. M CT\

1/5

I -

:
^ ^ ^

3C & o

«^

.S 5

w)

^ -^ I -s 1 "2 ? •§

•n2!6 i- 1I 3

ISl I

2

^8

•2^ cs

^e ^ o

•^ ^ 2

c« too

1 §

P ^

1^

^^

1|

i-y

' 3 m t 3

-^?>

-^^ s

Sw

CN

Q OH

CO

a u

g

o

O Q

G O

i|

m4^ 11 lli it

II

1it SDI

I

I

-« 73

21

«s o

..I 6 ^ —

r

«J

+ .2

on O

111!

.2? b

^ ^ '^

>

ill

P

II ^= 1

•

m

!

j « &s *

Q

o

i

.2 2 ^

•E V. 2

jlillli I

li I'

liiii ! |

iii§

f

Iff

IIS




B

-S -§ 2^ S § 2 ^ 3 3 B S ^ ^ § 3 -2 g PC

^ S

c 1>5 •J3 S irrn o

^

O Q

II

s ffi

o c

II li 6

13 &

c«^

III S .« =1 a 00 o

^ i 6

-o

11

2^

1 1a ^

11 g§ 4^ t S ao O -^

o

o u

i!^ :z;

:
n r»^ m in fn

^

?Q I ^ &

_

1^ < :i

^ • • S ' ^

^1^

11

^ 1 2^^ I

g^

8 %

u

Chromophoric DOM in the Coastal Environment

523

The parameter, S, defines the spectral dependence of the CDOM absorption coefficient, and thus provides information about the "nature" of the CDOM chromophores. Unfortunately, in the past, researchers have not determined this parameter in a consistent fashion. Most investigators have calculated S through a linear least squares regression of the log-transformed absorption data (Table I; Fig.3B;Bricaude?(3/., 1981;Blough^ra/., 1993; Green and Blough, 1994;Nelson and Guarda, 1995; Nelson et al, 1998). However, as pointed out by a number of workers (Stedmon et al, 2000; Boss and Twardowski, private communication), fitting the data to an exponential form using a nonlinear least squares fitting routine represents a better approach, owing to the relatively greater weighting given to the higher, and better measured, absorption values at short wavelengths. In contrast, a linear fit to log-transformed data enhances the relative weights of the low absorption values at long wavelength, and thus biases the value of the slope downward (Fig. 3). This problem becomes particularly acute when workers attempt to fit data over spectral ranges where the absorption is close to the detection limit of the instrument. In contrast, the use of the nonlinear least squares regression allows the complete spectral range of the data to be fit (e.g., from 290 to 700 nm), thus foregoing the indiscriminate use of different spectral ranges to acquire S. A goodness-of-fit statistic, such as x ^» can be reported and used to evaluate whether a simple exponential model provides a sufficient description of the data, or requires the use of a more complicated model such as the sum of two exponentials. However, if the residuals from the single exponential fit fall within the photometric accuracy of the instrument (~0.046-0.115 m~^), the use of a more complicated model is difficult to justify (Fig. 3A, inset). In the past, we have employed a linear least squares regression to obtain S over the range from 290 nm to the wavelength where the detection limit of absorption is reached. A recent analysis in this laboratory has shown that the S values acquired in this fashion are biased toward lower values than those obtained using the nonlinear fitting by ~0.0023 nm~^ (Figs. 3 and 4). However, the values of S obtained in these two ways are well correlated (r^ = 0.857), with the relationship exhibiting a slope very close to one (Fig. 4). These results indicate that although the values of S acquired by the linear least squares analysis may slightly underestimate 5, observed changes in S will be of the same magnitude. Figures 3 A and 3C further illustrate that CDOM absorption spectra measured for waters having substantially different levels of absorption are very well described by an exponential function within the photometric accuracy of the measurements. Because many workers in the past have acquired S using different approaches (linear vs nonlinear) or over different spectral ranges or series of ranges, comparisons among studies can be difficult. Thus, we have annotated the spectral data presented in Table I with information on the method of measurement. S varies with the source of the CDOM (Table I), but also can be altered through the biological and chemical processing of a source material. Values of S for humic

Blough and Del Vecchio

524

0.030 H 31-33

0.025 H ^y!^ 20-31

^

0.020 H

0.015 H NLF = 0.98315 * LF + 0.00229 r 2 = 0.8569

0.010 0.010

T

0.015

0.020

T^

0.025

0.030

S (nm-^) (LF) Figure 4 Values of 5 obtained for CDOM absorption spectra for waters in the MAB during September 1998 using a nonlinear least-squares fit to an exponential curve (NLF) and a linear least-squares fit to a semilog plot (LF). The samples were collected along a transect from the Delaware River to ~375 km offshore; theflcDOM(355)ranged from 4 to ~0.1 m " ^

substances and for CDOM from a wide variety of sources range from as low as ^0.01 nm~^ for terrestrial humic acids to as high as 0.02-0.03 nm~^ for CDOM in oligotrophic seawaters (Table I; Fig. 4). For humic substances, the relationship between S and the "molecular" properties of these materials can be summarized as follows (Blough and Green, 1995): (1) S is larger for fulvic acids than for humic acids; (2) S increases with decreasing molecular weight; (3) S increases with decreasing aromatic content. Specific absorption coefficients, a(Xy, for humic substances, obtained by normalizing aiX) to the organic carbon concentration, C, a(X)*[L (mg org. C)"* m"^] = a(X)/C,

[3]

increase with increasing aromatic content (Chin et al, 1994). These results are consistent with the view that humic substances having higher aromatic content, generally those with higher molecular weights, exhibit lower values of S. These lower values arise from enhanced absorption at longer wavelengths (lower energies). This may result from the presence of a distinct suite of chromophores having extended aromatic systems that absorb at lower energies (longer wavelengths). Alternatively, the increased absorption at longer wavelengths could arise from intramolecular charge transfer transitions between (similar) chromophores due to their greater numbers (and thus higher aromaticity) (Power and Langford, 1988; Blough and Green, 1995). Similarly, the increase in a{Xy with increasing aromatic content can be explained as due to a higher percentage of light-absorbing aromatic structures or to an increase in the number of charge transfer interactions.

525

Chromophoric DOM in the Coastal Environment

A number of recent studies have found that the values of S art larger for offshore seawaters (>0.02 nm~^) than for coastal waters influenced by river input (0.013-0.018 nm-i) or for most fresh waters (Table I; Fig. 4; Brown, 1977; Carder et al, 1989; Blough et al, 1993, Green and Blough, 1994; Nelson and Guarda, 1995; Nelson et al, 1998). Despite a few reported exceptions (Stedmon et al, 2000), S is usually observed to increase with decreasing absorption and increasing salinity during transit of the terrestrial CDOM to offshore waters (Fig. 4; Blough et al, 1993), suggesting that this material is being altered or replaced with a marine form. During the summertime in the MAB, the dependence of CDOM absorption on salinity reveals evidence of a significant CDOM sink in surface waters under conditions in which S also increases (Figs. 5 and 6A; Vodacek et al, 1997), consistent with the view that the terrestrial CDOM is being altered (Sections III and IV below). Based on the properties of the humic substances, this increase in S suggests a loss of aromaticity and a decrease in the average molecular weight of the CDOM. Recent laboratory and fieldwork indicates that photochemical bleaching can produce the same effects (Vodacek et al, 1991 \ Schmitte-Kopplin et al, 1998; Moran et al, 2000) and may account in part for the observed gradient in S from in- to offshore waters. However, the in situ production of CDOM having higher

6 **v, ^ \

4 2

ID in c Li_

0 6

1.5

A 08/93, 20-35 m —* 11/93 • . . 03/94

-J0.5p,

•A

. ,

_ 1

1

^^^fW^a^^ ,

,

,

^i \ * |

4

,

,

,

,

,

,

,

1

,

,

,

1

A 08/93, 20-35 m

0 ^ =lT.5^

r-=

000O8/93

2

0.5

:B

0 26

• . ,° .^r*T'*.° .

28

30 32 salinity ( V J

34

.o°m.

36

Figure 5 Seasonal dependence of CDOM absorption coefficient (acu (355)) and fluorescence (excited at 355 nm, Fn(355)) on salinity for (A) surface waters from the Delaware River to the Sargasso Sea; (B) waters at the MAB shelf break. Arrows indicate the location of the shelf break defined as 200 m isobath. Reprinted with permission from Vodacek et al. (1997).

Blough and Del Vecchio

526

0.030 -

••

Vk

11

0.027 \ 0.024 '

C/5

o

0.021 \

. . . . . . . . ••• .'•« •° ' - ^ o °o 1

0.018 0.015 J

1

•

•

MAB mixed layer MAB below mixed layer

o

o

460-

• • •

440=3

s

1 sc o

V

V

y

V

••

•

•

420 • 400 J 380 ' 360 1

• ^

340 '

o

• •

•

V^ o A •

1 1 1

A O

Orinoco River (ex 350 mn) Gironde Estuary (ex 313 mn) Continental shelf (ex 313 nm) Mediterranean Sea (ex 313 nm) West Rorida Shelf (ex 285-310 nm)

320 •

\ ^^•^^f^"^^

10

15

20

25

30

35

40

Salinity (ppt) Figure 6 (A) Dependence of CDOM spectral slope {S) on salinity for the MAB, in the mixed layer ( • ) and below the mixed layer (O). [The 5 has been calculated with a nonlinear least-squares fit (NFL) over the range 290-700 nm. Similar results have been reported by Blough et al (1993) for the Orinoco River using a linear least fit (LF)]. (B) Dependence of CDOM fluorescence emission maximum on salinity in the Orinoco River ( • ) (Ex. 350 nm) (data from Del Castillo et al, 1999); the Gironde Estuary (V) (Ex. 313 nm) (data from de Souza Sierra et al, 1994, de Souza Sierra et al, 1997); the Continental Shelf (•) (Ex. 313 nm) and the Mediterranean Sea (O) (Ex. 313 nm) (data from de Souza Sierra et al, 1997); the West Horida Shelf (A) (Ex. 285-310 nm) (data from Del Castillo et al, 2000).

values of S presumably could also play a role, especially in oligotrophic waters far from the influence of terrestrial sources (Nelson et al, 1998).

B. FLUORESCENCE Fluorescence measurements of humic substances and CDOM have generally been more common than absorption measurements, due primarily to their greater

Chromophoric DOM in the Coastal Environment

517

sensitivity and simplicity (Donard et al, 1989; Chen and Bada, 1992; Green and Blough, 1994; De Souza Sierra et al, 1994, 1997). Fluorescence is far more amenable to continuous monitoring (Vodacek et al, 1995; Klinkhammer et al, 1997,2000; Chen, 1999; Guay et al, 1999) and remote measurement (Hoge et al, 1995b; Vodacek et al, 1995), and thus potentially allows for the high resolution mapping of CDOM distributions. However, fluorescence provides only an indirect measure of CDOM absorption, and its magnitude and spectral dependence are more sensitive to such factors as pH, ionic strength and the presence of quenchers. Steady-state fluorescence is not representative of the entire population of absorbing species within CDOM, nor of the entire population of emitting species; instead fluorescence spectra will tend to be dominated by those subpopulations exhibiting longer fluorescence lifetimes (Herbelin, 1994; Blough and Green, 1995; Lakowicz, 1983). Thus, to estimate CDOM absorption from fluorescence measurements, a linear relationship between fluorescence and absorption must be established empirically for the geographical region of interest (Section ILB.2). Spectrofluorometry can also be used to acquire additional information about the possible sources and nature of the CDOM through the collection of excitationemission matrix spectra and synchronous scan spectra (Cabaniss and Shuman, 1987; Coble et al, 1990; Green, 1992; Green and Blough, 1994; Blough and Green, 1995; Pullin and Cabaniss, 1997; Section II.B.l), keeping once again in mind that emission spectra tend to be dominated by the longer-lived fluorescent component(s). 1. Fluorescence Excitation and Emission Spectra, Excitation-Emission Matrix Spectra, and Synchronous Spectra The excitation and emission spectra of humic substances and CDOM are very broad and unstructured, with the maxima in the excitation and emission spectra usually falling between 300 and 400 nm and 400 and 500 nm, respectively. The emission maximum shifts continuously to the red and lowers in intensity with increasing excitation wavelength, suggesting the presence of numerous absorbing and emitting centers. Three-dimensional excitation, emission matrix spectra (EEMS) have thus been employed in an attempt to distinguish source-dependent variations in the CDOM. EEMS are obtained by acquiring emission spectra at a series of successively longer excitation wavelengths. These emission spectra are concatenated to generate a plot in whichfluorescenceintensity is displayed as a function of the excitation and emission wavelengths. In contrast, synchronous scan spectra are obtained by scanning the excitation and emission wavelengths simultaneously at a fixed wavelength difference and thus represent a "slice" through the EEMS (Blough and Green, 1995). Although slower to collect, EEMS provide a more complete picture of the CDOM emission properties and can often be used to discriminate among different classes

528

Blough and Del Vecchio Table II Excitation and Emission Maxima for Classes of CDOM Fluorophores Identified by EEMS Type of fluorophore Protein-like Tyrosine Tryptophan Humic-like UV-humic UV-humic Unknown Visible-marine humic Intermediate marine-terrestrial Visible-terrestrial humic Chlorophyll-like Chlorophyll

^ex'^em

Labeled

230, 275/305 230, 275/340

B T

230/430 260/400-460 280/370 290-310/370-410 310/412 320-360/420^60

A N M Intermediate C-M C

398/660

P

Note. Modified from Coble et al, (1998). Table 2, p. 2208.

of fluorophores based on their excitation/emission wavelength maxima, as well as to follow changes incurred by the biological or physical processing of a material. Through the use of EEMS, several broad classes of emitting species have been identified in natural waters, namely the "protein-like," the "humic-like" (CDOM), and the "chlorophyll-like"(Table II; Coble et al, 1990,1998; Mopper and Schultz, 1993; Coble, 1996). The humic-like class has excitation/emission maxima in the range 320-360/420-460 nm, respectively. Offshore waters exhibit excitation and emission maxima for this class that are blue-shifted by ~25 and ^5-30 nm, respectively, relative to coastal and estuarine waters (Fig. 6B; De Souza Sierra etal, 1994, 1997; Coble, 1996). Consistent with the increase in S observed in the absorption spectra (Fig. 6A), these blue-shifts point to a preferential loss of longer-wavelength, lower energy transitions and a decrease in aromaticity, produced possibly by the photochemical and/or microbial processing of the terrestrial CDOM, its mixing or replacement with a less aromatic marine form, or some combination of these effects. The "protein-like" class exhibits excitation maxima at 220 and 270 nm with emission maxima at 305 nm (similar to tyrosine) and at 345 nm (similar to tryptophan). As reported by Mopper and Schultz (1993), the intensity ratio of these excitation peaks is similar to that of proteins reported by Lakowicz (1983), suggesting that these amino acids, either free or within proteins, contribute to the fluorescence signal of some natural waters. These signals have been observed in surface water (Coble et al, 1990; Mopper and Schultz, 1993) and in porewaters (Coble, 1996).

Chromophoric DOM in the Coastal Environment

529

2. Fluorescence Quantum Yield and Fluorescence/Absorption Relation The fluorescence quantum yield is a well-defined photophysical parameter representing the ratio (or percentage) of photons emitted to those absorbed. This parameter can be used to compare the fluorescence efficiencies of CDOM from different locales, thus providing another tool for characterizing the "nature" of the CDOM. A constant quantum yield indicates that the fluorescence intensity of a material will be directly proportional to its absorbance at the excitation wavelength (in the absence of inner filtering; Lakowicz, 1999); the higher the quantum yield, the higher the fluorescence per unit absorbance. Thus fluorescence measurements can also be used to acquire CDOM absorption coefficients, if a linear relationship between a well-calibrated fluorescence signal and absorption can be established (Hoge et al, 1993; Green and Blough, 1994). In earlier work, Green and Blough (1994) showed that the quantum yield for fluorescence obtained with 355 nm excitation (^355) was relatively constant for different humic substances and for CDOM from significantly different geographical areas. 0355 varied by about a factor of five, with the CDOM of most waters having a ^355 of about 1% (see also Vodacek et al, 1995, 1997). Over this range, the highest yields were observed for humic substances isolated from deep marine waters (2.1%), whereas the lowest were observed for terrestrial humic acids and some river waters (~0.4%). The relative invariance of these yields across differing oceanic environments is also mirrored in the linear relationships between CDOM fluorescence and absorption that have been observed by numerous investigators over the past 10 years (Table I; Ferrari and Tassan, 1991; Hoge et al, 1993; Nieke et al, 1991 \ Vodacek et al, 1997; Ferrari and Dowell, 1998; Seritti et al, 1998; Ferrari, 2000). The largest difference in the ratio of absorption to fluorescence obtained from the slope of these relationships (m in Table I) is about a factor of 3 and is observed between the eastern and western coasts of the North Atlantic Ocean (Table I; Section III). However, within a given geographical area, the variation in the ratio is much less, generally no more than ^15-30%. Thus, these fluorescence/absorption relationships have been employed to acquire CDOM absorption coefficients from continuous measurements of m situ fluorescence, as well as from fluorescence measurements acquired by aircraft using NASA's Airborne Oceanographic Lidar (Hoge ^r a/., 1995b; Vodacek ^r a/., 1995, 1997).

3. Relationship of Absorption, Fluorescence to DOC Despite the fact that a significant fraction of the dissolved organic carbon (DOC) is not associated with CDOM, a number of workers have observed correlations between CDOM absorption (orfluorescence)and DOC concentration in the coastal environment (Fig. 7; Laane and Koole, 1982; Vodacek et al, 1995, 1997; Chen,

Blough and Del Vecchio

530 —!

^

1

T

F

,

+ Nov, Mar, Apr mixed layer o Aug below mixed layer • Aug mixed loyer 4-

0

O O

_j 1.5 -I

>/^

r in

r

•

•

in

y ^

••

-1

60

80

100 120 DOC (/xM C)

0.5

1

1

1 — .

140

.

o (d

0 160

1—1

Figure 7 Seasonal dependence of CDOM fluorescence (Fn(355)) (left axis) and absorption (acM(355)) (right axis) on DOC in the MAB. The regression line (slope of 0.0157 and intercept of 67 /xM) refers to data from November, March, and April. Reprinted with permission from Vodacek et al, 1997.

1999; Klinkhammer et al, 2000). These correlations usually exhibit a substantial positive intercept on the DOC axis (Fig. 7), implying that the oceanic end-member contains predominantly nonabsorbing DOC and little CDOM. The large intercept and steep slope of this correlation results from the very different content of CDOM and DOC in the freshwater and oceanic end-members; while DOC decreases only by about a factor of four between fresh waters and the surface ocean (~300 fiM vs --70 /xM), CDOM absorption decreases by factors ranging from 40 to 200 (Table I; Fig. 8). This relationship appears to arise primarily from the quasi-conservative mixing of both CDOM and DOC within these regions (e.g., Mantoura and Woodward, 1983), and not through a covariation in the rates of their in situ production and consumption. Under conditions in which CDOM is consumed photochemically (Figs. 5 and 6), this relationship is altered (Fig. 7; Section IV). These results are unlike those observed for the open ocean (Nelson et ai, 1998) or for coastal regions not strongly influenced by river inputs (Ferrari, 2000). In these and other regions such as the northwest Florida shelf (Del Castillo et al, 2000), no (or weaker) correlations have been observed between DOC and CDOM, due to an uncoupling between the sources and sinks of the CDOM and the nonabsorbing DOM in this environment.

Chromophoric DOM in the Coastal Environment

531

Salinity (ppt) Figure 8 Dependence of acDOM(355) on salinity for different geographical areas: (A) Orinoco River on 9/27/88 ( • ) and on 10/9/88 (O) (data from Blough et al, 1993); Gulf of Paria (T) (data from Blough et al, 1993); MAB on April 1994 (V) (data from Vodacek et al, 1997); SAB on August 1992 ( • ) and on April 1993 (D) (data from Nelson and Guarda, 1995); Apalachicola River (•) and Suwanee River (O) (data from Del Castillo et al, 2000); St. Lawrence Estuary (A) (data from Nieke et al, 1997); Amazon River (A) (data from Green and Blough, 1994). (B) North and Baltic Sea ( • ) (data from H0jersley et al, 1996); Baltic Sea (O) (data from Stedmon et al, 2000); Eastern Atlantic Ocean (T) (data from Ferrari, 2000); Tyrrhenian Sea and Amo River (V) (data from Seritti et al, 1998). (Inset) Same Eastern Atlantic Ocean data ( • ) with expanded axis (data from Ferrari (2000)). The «CDOM(355) has been recalculated using theflcDQM(^)and the spectral slope reported for each study.

532

Blough and Del Vecchio

III. DISTRIBUTION Field measurements of CDOM optical properties have increased enormously over the past decade. However, because workers have often used different experimental protocols, a direct comparison of the optical properties of CDOM from different waters is often difficult. Nevertheless, we have attempted to compile in Table I a representative list of CDOM optical data obtained in studies over the past ~10 years, and where possible, have converted these data to a common set of parameters so that comparisons could be made (Fig. 8). Although many blackwater rivers such as the Tamiami (southwest Rorida; Green and Blough, 1994), the Surumoni (South America; Battin, 1998), and the Satilla (Georgia; Moran et ai, 2000) can exhibit very high values of «CDOM(355) (> 30 m~ ^), the mouths of the larger rivers and estuaries generally have much lower values, on the order of 5 to 15 m"^ (Fig. 8; Table I). With one notable exception (the outflow of the Orinoco River), acDOM(355) is inversely related to salinity, and for many estuaries and coastal waters appears to behave conservatively, although not in all cases nor in all seasons (Figs. 5 and 8). Nonlinear mixing curves can arise from the in situ production or loss of the CDOM, from the conservative mixing of three or more water masses containing different acDOM(355) end members or from some combination of these factors. Depending on the locale, it is often difficult to distinguish between the in situ production and consumption of CDOM versus the mixing of different water masses. As one example, the dependence of acDOM(355) on salinity for the Orinoco River implies the presence of a large source in the region of the outflow (Fig. 8; Blough et al, 1993). Based on additional evidence, this CDOM does not appear to arise from in situ phytoplankton production, from a sediment source nor through release from particulate matter at higher salinities. An alternative, but as yet untested possibility, is that there are other water masses containing very high levels of CDOM intruding into the Orinoco delta. The nonlinear dependence of CDOM absorption on salinity observed by H0jerslev et al (1996) in the North Sea-Baltic Sea transition zone (Fig. 8) has been interpreted to result from the quasi-conservative mixing of (terrestrial) CDOM from three water masses: (1) North Sea water (high salinity of 35, low CDOM— «CDOM(355) = 0.099 m~^); (2) Baltic Sea water (low salinity of 8, intermediate CDOM—acDOM(355) = 1.36 m"^); (3) German Bight/southern North Sea water (intermediate, high salinity of 31, high CDOM—nflcDOM(355) = 2.13 m"^). At salinities lower than 8, the CDOM increases even further due to the contribution of freshwater inputs from the Bothnian Bay (Fig. 8). Based on their measurements and historical data, these workers concluded that the long-term average concentration of CDOM had not changed significantly over the past 40 years in the Baltic, the North Sea or the Atlantic, suggesting that the (terrestrial) source is in balance

Chromophoric DOM in the Coastal Environment

533

with a loss pathway. They further concluded that there were only minor seasonal variations in the CDOM levels. However, depending on the spatial and temporal scales examined, workers have come to quite different conclusions concerning the dynamics and seasonal distributions of CDOM. As an example, in the southern Baltic Sea (Ferrari and Dowell, 1998; Kowalczuk, 1999), the magnitude of CDOM absorption was found to vary seasonally depending on the river input to the near-shore bay waters and was inversely related to salinity. In contrast, no correlation of absorption with salinity was observed in offshore waters, possibly due to the numerous river sources contributing to this region. However, a significant correlation was also observed between CDOM absorption and chlorophyll (Chi) concentration in the offshore waters (Kowalczuk, 1999), suggesting that the in situ formation of CDOM from phytoplankton was also playing a role. In general, coastal areas subject to high river inputs exhibit high levels of CDOM absorption, with the magnitude of the absorption depending on the river end member(s) contributing to the region and the seasonal river flow(s). Absorption is inversely related to salinity and exhibits conservative mixing behavior (Fig. 8) in the absence of significant in situ sources and sinks or mixing between multiple water masses. The geographical area impacted varies seasonally, depending on the magnitude of the river flow(s). In contrast, coastal margins not affected by river inputs generally show low values of acDOM(355) ( nd nd 4.5^'

486

1.8^'

5.0/

5,780 1,100

470 145 215 120

5,840

3.5 1.2

1,300

1.0

152 10 38 75 11

29.8

3,409

1.9 2.0 20 0.7 211

433 555 505

13.4 14.5 11.7

0.27

3.0 0.54

2.99 2.50 2.44 0.358 0.75

POC (xlO^t/year)

TSS (xlO^t/year)

Volume (km^/year)

27.3 47.5 34.2

25.4

13.0« 2.0^ 1.3^ 0.1^ nd nd 24.1

1.95

925

0.015 0.35

222

nd nd

nd 1.1^

nd nd

0.79

666

nd

nd

nd

14.7

(Continues)

582

Gustave Cauwet Table I (Continued) River

Irrawady Ganges + Brahmaputra Indus Estimate of total continental Europe Wolga Don Dniepr Danube Po Tiber Rhone Loire Seine Garonne Rhine Elbe Vistula Northern Dvina Pechora Estimate of total continental Estimate of total (excl. Australia)

POC (xlO^t/year)

Area (xlO^km^)

Volume (km"^/year)

TSS (xlO^t/year)

DOC (xlO^t/year)

0.43 1.48

428 971

265 1,670

nd 3.6'

nd 32'

1.17 44.1

238 12,205

100 11,172

0.75' 94

nd 128

1.46 0.43 0.53 0.82 0.067 0.017 0.099 0.121 0.079 0.085 0.224 0.146 0.199 0.365 0.330 10

243 29.3 52.3 198 46.4 7.2 59.9 27.0 15.8 21.4 69.4 23.7 34.7 112 128 2,826

27.4 6.4 2.12 83 9.0 nd 11 7.8 3.54 1.3 3.4 0.84 nd 1.54 1.44 158

nd nd nd 0.59^" 0.12" nd 0.18^ O.llP nd 0.075P nd nd nd nd

nd nd nd 0.356'" 0.14" nd 0.09^ 0.04P nd 0.035P 0.37^? nd nd nd nd

126

35,300

4,625

250

176

Note. Modified with permission from Degens et al, 1991. ^Richey(1991). ^Depetris and Paolini (1991). ^Telangera/. (1991). ^Prahl and Coble (1994). ^Hopkinson^ra/. (1998). /Leenheer (1982). ^Martins and Probst (1991). ^Olsson and Anderson (1997). ^Cauwet and Sidorov (1996). ^Cauwet and Mackenzie (1993). ^Cauwet, impubUshed data. 'Spitzy and Leenheer (1991). ^Cauwet et al. (in press). "PettineeM/. (1998). ^Cauwet ^r a/. (1990). ^'Kempeera/. (1991). '^Eismaera/. (1982). ^Laraera/. (1998).

583

DOM in the Coastal Zone

Table II DOC and Water Export from Major Morphoclimatic Zones Based on Discharge and Typical DOC Concentration Data After Meybeck (1988) and Spitzy and Leenheer (1991). Water discharge

DOC export

Morphoclimatic zone

Typical DOC concentration (mg/L)

km^/year

% of total

Tundra Taiga Temperate Wet tropic Dry tropic Semiarid

2 7 4 8 3 1

1,122 4,376 10,285 19,186 2,169 262

3 11.7 27.5 51.3 5.8 0.7

2.2 30.6 41.1 153.5 6.5 0.3

1 13 17.6 65.6 2.8 0.1

37,400

100

243.2

100

Total

10^ t/year

% of total

Amazon River, the most recent data (Hedges et al, 1994) shows 300-400 /xM DOC in the lower main stem of the river, with high values in some black water tributaries (Rio Negro, 800 JJM). For Arctic Siberian rivers, we now have reliable DOC measurements (Cauwet and Sidorov, 1996; Lara et aL, 1998) that can replace the old TOC data used in the Spitzy and Leenheer budget. The mean DOC concentration in the lower Lena River is probably close to 600 /JM from June to September, a period that represents about 85% of the total annual discharge. The resulting annual DOC flux is about 3.6 x 10^ t C. For a more detailed study of carbon inputs to the Arctic Ocean, see Olsson (Olsson and Anderson, 1997; Anderson, Chapter 14). For the main Chinese rivers (Yangtze, Huanghe, and Pearl Rivers) more data have been recently collected, and a large correction to estimated fluxes has to be done. The mean DOC concentration for the Huanghe River (Yellow River) was greatly overestimated at 1000 /xM and is probably closer to 250 /xM (Zhang et aL, 1992; Cauwet and Mackenzie, 1993). In addition, the discharge of the Yellow River has decreased over the past 20 years due to damming and irrigation. Flow is now probably less than 600 m^ s~^ while it was estimated around 1300 m^ s~^ in previous studies. Consequently, the annual DOC flux of the Huanghe River previously proposed (0.54 x 10^ t year"^) has to be divided by almost 8 (0.076 x 10^ t year~^). DOC concentrations were also overestimated for the Yangtze River. It was recently estimated in the range 250-400 /xM (Cai et aL, 1992; Cauwet and Mackenzie, 1993), much less than the 1100 /xM (13.4 mg C L~^) cited by Spitzy and Leenheer (1991). For the Pearl River there are only infrequent and recent DOC data (Cauwet, unpublished). The estimated DOC concentration is around 600-800 /xM C. These values were measured on only a few samples collected in an estuarine environment and could be influenced by human activity, leading to an overestimation of the natural concentration in

584

Gustave Cauwet

the river, but contributing partly to the total input to the coastal sea. Some recent studies on the Mississippi River (Gardner et al, 1996; Kelley et al, 1998) also give estimates of DOC concentration but the new data do not change significantly the earlier carbon budget. If we reassess the global DOC input from rivers to oceans, we need to modify the amount (0.11-0.25 10^ t year"^) estimated by Degens and Ittekkot (1985). An estimate made by Degens et al. (1991b) and reproduced by Meybeck (1993) proposed a total DOC input of 0.2 x 10^ t C year~^ but without European rivers and without any data on the Ganges, Brahmaputra, Irrawady (no data exist for that large river), Mekong, Lena, Ob, and lenissei in Asia, Zambezi and Senegal, in Africa, or Rio Magdalena in South America. If one considers that these rivers contribute about 14% of total freshwater discharge to the global ocean and include some of the large rivers having the highest DOC concentrations (the Siberian rivers), we can reasonably estimate that the total DOC input to the global ocean is about 0.25 x 10^ t year"^ or 0.25 x lO^^g C year"^ (0.25 Gt C year"^). This estimation is comparable to that made recently by Hedges et al. (1997). This annual input represents only about 0.0004 times the DOC content of the ocean (estimated at 685 Gt C; Hansell and Carlson, 1998a), but sufficient alone to sustain the 6000-year estimated turnover of DOC of the ocean (WiUiams and Druffel, 1987). This riverine input remains low compared to the annual production of the whole ocean (about 50 Gt C year~^).

B . BlODEGRADABILITY OF RiVERINE D O M Though the yearly DOC discharged by rivers represent only 0.03% of total marine DOC pool, the impact of DOC inputs on the coastal zone is far from negligible. Furthermore, it is unknown how much is rapidly degraded or persists in the marine environment. Hedges et al. (1997) have documented the recent studies on the topic and discussed the fate of terrestrial organic matter in the ocean. They compiled all data (bulk composition, isotopic characteristics, molecular tracers approach) that were used to differentiate terrestrial from marine organic matter, mostly on particulate matter and coastal sediments but sometimes on the dissolved fraction. They concluded that either our global budgets and distribution estimates are greatly in error, or both dissolved and particulate organic matter of terrestrial origin suffer rapid and remarkably extensive remineralization at sea. The first approach utilized to evaluate the possible biodegradability of dissolved or particulate organic matter was the analysis of the probable most degradable fractions: proteins and carbohydrates. In an early work, Ittekkot and coworkers collected all data on carbohydrates and amino acids in the world's rivers (Ittekkot et al., 1982). These authors estimated that carbohydrates and amino acids mainly represent the aquatic life and present a high degradable character, in contrast to

585

DOM in the Coastal Zone

4

3] (0

2

O

1 3

4

5

9

6

10

11

Months

1200

900 800 700 S 3 O Q

1000 800

600 500 400 300 200

5" - =^

1000 Da) while this HMW portion is only 30% in surface ocean and 22% in deep ocean (Hedges et al, 1994). Ogawa (1999) also mentioned that the structure of HMW DOM is similar to biomolecules while the LMW fraction seems more complex in terms of structure and poorly determined. Looking at the literature, it is rather difficult to have a precise idea of the nature of dissolved organic matter discharged by rivers. There are several reasons for that, the main one being that information is sometimes for dissolved, sometimes for particulate fractions, and sometimes for the total organic matter. For instance, cellulose (and relative structures) and lignin are major constituents of terrestrial plants, but these compounds are not known to be very soluble and they are mostly transported as detritus. To get soluble fractions there must be degradation by microorganisms in the soils or in the water. Consequently, the DOM from soils is discharged when soil leaching occurs, generally during rainy seasons and high discharge periods. This is what occurs with the increase in DOC concentration during floods in many rivers. However, it is not that simple; heavy rains bringing large quantities of water that dilute the leachate, and what is often observed is an increase in DOC at the beginning of the flood and a decrease, due to dilution, in the second part of the flood (see Fig. 1). In fact, the DOM in rivers is mainly composed of humic matter, bringing with it the degraded characteristics of lignin

587

DOM in the Coastal Zone 1000 g* 800 3 600

g

Q

400 ]

200 -{ 0 1

2

3

Q(10^xmV) ^^ 700 O) 600 3 500 ^ 400 ] o 300 -]

2 200

c 100 E 0 < 0 Q{10^xmV')

3

(0

1200 n 1000 800 -

&

600 -

o

400 -

•o

•s (0 O

200 n U i

0

•

t 1

2

3

4

Q(10^xmV) Figure 2 Relationship between water discharge (Q) and DOC (a), amino acid (b), and carbohydrate (c) concentrations in the River Ganges.

and cellulose, but also a minor fraction originating in the riverine production. This does not exclude direct human influence, particularly in estuaries surrounded by regions of dense population, where organic matter issued from pollution can be an important and rapidly mobilizable fraction. Nevertheless, most authors point out the difficulty in giving a precise estimation of the real flux of DOM issued from

588

Gustave Cauwet

rivers, even with high frequency sampHng. A perfectly representative sampling of flood events is almost impossible in terms of the relationship between discharge and concentration. This is essentially due to the high variability observed in the river water at a space scale (the sampling strategy can be improved for large rivers) and at different time scales (daily, annual, and interannual). A good example can be cited with a 2-year (13 cruises) survey of the Chesapeake Bay (Fisher et al, 1998), where DOC concentration varied in the river from 129 to 316 /xM (mean value 232) and was measured in the estuarine zone from 144 to 374 jiM. No clear relation can be shown between DOC concentrations in the river or the estuary with the Susquehanna River discharge; there is little seasonal variation. Such a detailed survey aptly describes the processes occurring in the mixing zone, especially identifying the internal sources. Nevertheless, the authors point out that the interpretation of mixing diagrams must be done with caution, due to the variability of the end-members that can create a nonlinear behavior under conservative mixing. This was already mentioned by other authors (Loder and Reichart, 1981; Cifuentes et al, 1990) in the past and recently. In conclusion, most of the recently published works in estuaries and the coastal zone show that the most important source of dynamics for carbon is probably linked to the equilibrium (or disequilibrium) existing between production and consumption of organic matter in the mixing zone. Now that we are more in confidence with most analytical data, the challenge is to identify the mechanisms transforming or oxidizing the allochthonous riverine DOM to explain the general observation that terrestrial organic matter is in low concentration in seawater and marine sediments (Hedges et al, 1997). The study of estuarine processes can certainly help to find the pathways of DOM from rivers to the sea, including the identification of internal sources and consumption mechanisms.

III. ESTUARINE PROCESSES A. PHYSICAL PROCESSES Due to high concentrations of DOM and particles and to the rapid change in salinity occurring in estuaries, a strong influence is expected of chemical-physical processes like adsorption, desorption, aggregation, flocculation, and deflocculation on the behavior of DOM. The earliest work describing flocculation processes in estuaries is probably that of Sholkovitz and coworkers (Sholkovitz, 1976; Sholkovitz et al, 1978). In these works, the authors described an important flocculation of dissolved humic substances (60-80% of the humic material was removed along the salinity gradient) and suggested that this humic fraction represented only 3-6% of riverine DOM. Nourredin and Courtot (1989) estimated that humic matter represents up to 60% of DOM and behave conservatively. In the estuarine flocculation

DOM in the Coastal Zone

589

literature it appears that each author has seen what he expected, relative to his own field. Specialists in clay focused on the particle size and the clay composition to explain fast deposition of sediment in estuarine environments (ICranck, 1981). Others pointed out that iron also participates in the flocculation and probably is the initiator of the mechanism (Sholkovitz et al, 1978; Fox, 1983, 1984). Others suggested that temperature is possibly an important factor (Droppo et al, 1998). Most of the time, especially when the results came from laboratory experiments, the authors did not take into account the water dynamics and the probable role it played. In fact, it seems that flocculation and deflocculation can occur at different stages of the estuarine mixing, in relation to mixing process (mixed or stratified estuaries) and the level of energy of turbulence (Eisma, 1986). It was sometimes shown that besides the turbidity maximum due to resuspension another turbidity peak could be attributed to a flocculation mechanism (Biggs et al, 1983). In well-mixed estuaries, it is difficult to differentiate these mechanisms; the rapid mixing and the presence of a turbidity maximum due to sediment resuspension completely mask the minor processes. In partly mixed estuaries, it is sometimes possible to observe this coupled mechanism showing exchanges between particulate and dissolved organic matter (Cauwet, 1985; Cauwet and Meybeck, 1987); but this occurs mostly in the freshwater portion of the estuary under tidal influence, as a consequence of variation in turbulence, not because of salinity increase. In well stratified estuaries (generally microtidal estuaries) it is easier to observe some flocculation processes occurring at the surface of the salt wedge (Cauwet, 1991; Lipiatou et al, 1991; Lipiatou and Saliot, 1992). The main reason is more probably the increase of the residence time of riverine colloids with low turbulence at the interface, rather than a drastic change in environmental conditions. It was observed that some molecular characteristics of riverine dissolved organic matter that are absent in particles brought by the river and in the river plume are again observed in particles found in the bottom nepheloid layer below the plume (Lipiatou et al, 1991). This finding suggests that some flocculation mechanism, involving dissolved or colloidal organic matter, occurs at the freshwater-saltwater interface and the newly formed floes are transported along the salt wedge toward bottom water in the estuary. One prominent character of riverine DOM that can influence the flocculation is the high concentration of colloids (Whitehouse et al, 1989; Filella and Buffle, 1993; Filella et al, 1993; Dai et al, 1995; Kim et al, 1995; Sempere and Cauwet, 1995; Gustafsson and Gschwend, 1997). Up to 50% of "dissolved" organic carbon (passing through GF/F glass fiber filters) is in the form of colloids and does not pass through 10-kDa membranes (Newman et al, 1994; Dai et al, 1995; Cauwet and Sidorov, 1996; Patel et al, 1999; Wells, Chapter 7). In many cases, the colloid concentration decreases rapidly with increasing salinity (Fig. 3). However, it has not been possible to demonstrate that a decrease in colloidal carbon corresponds quantitatively with a decrease in DOC or an increase in POC. Obviously, the exchange processes between the three fractions (true

590

Gustave Cauwet r-

80

7 5 Q 40 .•2^5

i

5^20 o

O

n 2 Salinity

Figure 3 Concentration of organic colloids along the salinity gradient at low salinity in the Lena River mixing zone.

dissolved, colloidal, and particulate) are more complex than sometimes supposed and the disappearance of colloids can enrich particulate matter or return back to dissolved phase. Fragile particulate flocculants can turn to colloidal or dissolved organic matter. These colloids, while accumulating at the freshwater-saltwater interface, have more time to coagulate and form aggregates. Several observations in estuarine environments have demonstrated that large flocculants (3-4 mm) are easily destroyed by turbulence while small flocculants (

• • o

^^^^Na

• 1

1

H 0


0 year~^) sediments in Fig. 6B. At the same time, though, while changes in a, Dj/Ds or kj; have similar effects on DOC pore-water concentrations in the surface mixed or burrowed layer, their effects on benthic DOC fluxes are quite different. As expected, increasing bioturbation or bioirrigation (at a constant kr value) leads to increases in the benthic DOC flux (Table I) in spite of observed decreases in the DOC pore-water gradient at the sediment-water interface (e.g., see Fig. 5). In contrast, enhanced DOC oxidation in mixed redox sediments (i.e., increasing k^ at fixed a or Dj/Ds values) decreases benthic DOC fluxes as a result of more effective utilization of DOC produced in situ during sediment remineralization. However, in this discussion I have assumed that both k^ and either a or Dj/Ds are affected by the occurrence of bioturbation or bioirrigation (the former as a result of associated changes in sediment redox conditions). Therefore, for a fixed input of reactive carbon to sediments (i.e., constant RQ and Roo values) these two opposing phenomena partially cancel one another out and lead to a slight enhancement in the benthic DOC flux as a result of their combined effects (Table I). In its present format, this model does not incorporate the effects of bioturbation or bioirrigation on the depth-dependence of HMW DOM production due to, e.g., mixing of reactive particulate matter deeper into the sediments (e.g., Aller, 1982; Smith et al, 1993; Hammond et al, 1996). Although such processes are likely to have additional effects on sediment DOC profiles, the results shown here indicate some of the ways that macrofaunal transport processes such as bioturbation or bioirrigation and accompanying changes in sediment redox conditions affect sediment DOC concentrations, cycling, and benthic fluxes.

C. CONTROLS ON DOC CONCENTRATIONS WITH DEPTH IN SuRFiciAL ( " S H A L L O W " ) S E D I M E N T S

DOC (and DON) accumulate with depth in sediment pore waters due to a slight imbalance between production and consumption (Burdige and Gardner, 1998;

Figure 6 Sensitivity plots examining average model-derived DOC concentrations in the upper 10 cm of sediment as a function of a (A), k^ (B and D), and D j / A (C). Note that kx = \,a =0 year~^ and Z>r/Ds = 1 define conditions in completely anoxic sediments (i.e., Fig. 4). All other parameters used in these calculations are defined in the captions to Figs. 4 and 5.

624

David J. Burdige Table I The Effects of Bioturbation, Bioirrigation, and Redox-Associated Changes in Sediment DOC Cycling on Model-Derived Benthic DOC Fluxes^ Benthic DOC flux^ ^r

a(year ^)

I>T/DS

mmol/m^/day

%of' "Anoxic''flux'^

HMW DOC flux'^ (% of DOC flux)

Bioturbation 1 1 10 25

0 0 0 0

1 3 3 3

4.28 6.22 5.27 4.86

100 146 123 114

60 67 79 86

Bioirrigation 1 1 10 25

0 60 60 60

1 1 1 1

4.28 6.49 5.83 5.06

100 152 136 118

60 58 65 75

^ All other parameters used in these model calculations are listed in the caption to Fig. 4. ^Benthic DOC fluxes were determined from the sum of the diffusive/bioturbation flux and the bioirrigation flux. Each of these fluxes was calculated individually for the HMW DOC and pLMW DOC fractions and then summed to yield the total DOCfluxesshown here. Diffusive/bioturbation benthic fluxes were determined using the equation -0Z>r(dC/dj:);c=oThe gradient at the sediment-water interface ([dC/djc];c=o) was estimated as AC/Ax using numerical results obtained with the BBS model. Here AC is the difference between the concentration in the overlying waters and at the first model grid point (at x = 0.05 cm), and AJC is therefore 0.05 cm. The bioirrigation flux was determined using the equation -/a(C-Co) dx (Boudreau, 1997). This integration was performed numerically (using trapezoidal approximations) over the depth of burrowed sediment using the results of the BBS model. ^Benthicfluxesof HMW DOC were calculated as discussed above for total DOC using calculated HMW DOC pore water profiles. They are reported here as a percentage of the benthic flux of total DOC shown in the fourth column. ^''Anoxic'' sediments are nonbioturbated and nonbioirrigated sediments (i.e., kr = 1, a = 0 year"^ Dj/Ds = 1). The results for anoxic sediments are shown in the first row of each section in this table. Alperin et aL, 1999). In surficial anoxic sediments these profiles generally approach asymptotic concentrations with depth (Fig. 1; again note that here surficial is defined as being less than ^ 1 m and is often on the order of the upper ~20-30 cm of sediment). This accumulation occurs predominantly in the form of refractory pLMW DOM (see Fig. 4), since both HMW DOM and mLMW DOM represent reactive DOM pools that turn over rapidly. As shown in Fig. 7 there is a positive relationship between maximum pore-water DOC concentrations ([DOC]oo) in anoxic surficial sediments and depth-integrated sediment carbon oxidation rates (Cox). Bioturbated/bioirrigated sediments do not appear to fall on the trend line for

Sediment Pore Waters

625

[m ( D)Calif. Borderland 10 \F T ( V ) Central Calif, margin [ [ L

u

1 SB

O mid-Atl shelf/slope break O Ches. Bay sta. S3 # Ches. Bay sta. M3 A ( A ) Other sites

•^

CLB

^^A

y ^

anoxic sediments

o

1k

s

0.3 t

1/

T «•^ J B^

I D ^

*

•^•n •

^

o o ^o__

sta. N

0.1 0.1

• y^%

"A \

bioturbated/bioirrigated sediments

10

100 -2 j - l x

Sediment Carbon Oxidation Rate (mmol m" d" ) Figure 7 The maximum DOC concentration in the upper ~20-30 cm of sediment versus the depthintegrated sediment carbon oxidation rate. Open symbols represent bioturbated/bioirrigated sediments while closed symbols represent more strict anoxic sediments. The two lines "through" these data sets are not meant to imply any functional relationships, but are simply presented here to show the different general trends in the data sets. Data sources: Chesapeake Bay sites M3 and S3—Burdige and Homstead (1994), Burdige and Zheng (1998), and Burdige (2001); California Borderlands and central California margin sites—Berelson et al. (1996) and Burdige et al (1999, unpubHshed data); midAtlantic-shelf/slope break (site WC4)—Burdige et al. (1996, 2000, unpublished data); Cape Lookout Bight, NC (CLB)—Martens et al. (1992) and Alperin et al. (1994); Skan Bay (SB), Alaska—Alperin et al. (1992); station N (see Fig. 2)—Bauer et al. (1995).

anoxic sediments, consistent with the discussion in the previous section regarding the role of macrofaunal processes in affecting pore-water DOC concentrations. At least two possibilities may explain these observation in anoxic sediments. The first is that a balance occurs at depth between DOC production (from sediment POM) and DOC consumption (Alperin et al, 1994; Burdige and Gardner, 1998). This explanation is implicitly incorporated into the ANS model since in Eq. [A-2] the concentration of pLMW DOC at depth (which is essentially [DOC]oo) equals the parameter Qs (= aR^/kp). Based on the model in Fig. 3 Qs is the steady-state concentration of pLMW DOC at depth that is achieved when consumption and production are balanced. If a and kp are roughly constant in these sediments (see Section II.D for further details), then Q^ is proportional to Roo and the observations in Fig. 7 are consistent with this explanation if /?oois positively correlated with Cox (which does not appear to be an unreasonable assumption). Interestingly, Alperin et al. (1999) used a very different modeling approach to examine sediment DOC

626

David]. Burdige

cycling and obtained the following proportionality: [DOCloc oc CoxZ*,

[3]

where z*is the ^-folding depth for sediment organic matter remineralization. Assuming that z* is roughly constant in all of these sediments this equation is also consistent with the observations in Fig. 7. A second explanation for asymptotic DOC concentrations with depth is that DOC production rates go to zero with depth and biotic or abiotic changes in the composition of the refractory pore water DOC pool (^pLMW DOC) continually decreases its reactivity. This would eventually lead to a situation in which pore water DOC found at depth is effectively nonreactive on early diagenetic time scales and is therefore selectively preserved. In this case, one might think of this DOC at depth much like one thinks of "inert" inorganic nutrient end products such as phosphate, ammonium, or X!C^2, which also show similar exponential-like profiles in anoxic sediments (e.g., Bemer, 1980). This analogy would predict that greater amounts of DOC accumulate with depth in sediment pore waters as rates of sediment carbon oxidation increase (e.g., Krom and Westrich, 1981), as is seen in Fig. 7. Additional insights into this possible explanation will be discussed in the next section. Finally, recent studies have shown that sorption of DOC to sediment particles can affect pore-water DOC concentrations (Hedges and Keil, 1995; Henrichs, 1995), and that pore water DOC concentrations may be "buffered" by reversibly sorbed DOC in equilibrium with the pore waters (Thimsen and Keil, 1998). While it is not inmiediately apparent how these processes could explain the results in Fig. 7 they could possibly affect DOC concentrations at depth depending on the intrinsic reactivity of pore water DOC and that which is adsorbed to sediment particles (Lee, 1994; Mayer, 1994b; Henrichs, 1995), the relative sizes of the pore water and sorbed DOC pools (Thimsen and Keil, 1998), and the extent to which sorption sites on particles deposited in a marine sediment are "saturated" by DOC-particle interactions in the water colunm. However, Alperin et al. (1999) recently concluded that the buffering of pore-water DOC concentrations by reversible sorption is not an important controlling factor in explaining pore-water DOC concentrations in these North Carolina continental slope sediments (see Section VI for further details).

D. PORE WATER DOC PROFILES IN DEEP SEDIMENT CORES In contrast to the numerous DOC profiles that have been collected in "shallow" surficial sediments (see Section I.A), far fewer studies have examined porewater DOC concentrations over larger depth and time scales (Starikova, 1970; Nissenbaum et al, 1972; MichaeHs et al, 1982; Chen et al, 1993; Alperin et al.

627

Sediment Pore Waters DOC (mM) 0

1

2

DOC (mM)

DOC (mM) 3

2

4

4

6

8

12

1

^

1 • x

1 L

•

I fit to entire data set

sta. M3, Chsapeake Bay (CH XVII, 8/96)

sta. C, North Carolina continental slope

\*

• •

!

•

fit to upper 115m

1 | •

ODP site 1082, Walvis Basin

Figure 8 Pore-water DOC concentrations versus depth from three contrasting anoxic marine sediments fit to Eq. [4]. The resulting best fit parameters are Hsted in Table II, along with the references to the sources of the data used in these calculations. Note the factor of 4 difference in concentration scales and factor of > 1000 difference in depth scales as one moves from the estuarine Chesapeake Bay sediments to the deep sediment (ODP) cores collected in Walvis Basin. The different symbols for the Chesapeake Bay plot (left) represent replicate cores collected on this date at this site. At site C on the North Carolina continental slope (middle) the data were fit to Eq. [4] starting at a sediment depth of 25 cm based on the observation that the upper portion of these sediments are extensively bioturbated (AlpQTinetaL, 1999). Therefore the term yx inEq. [4] was replaced here with y(x-25) and CQ (at 25 cm) was used as an adjustable fitting parameter. The resulting best value of Co was 1.42 mM versus the measured value of 1.50 mM. Finally, for the Walvis Basin sediments (right) both the entire data set and the upper 115 m of sediment were both fit to Eq. [4]. Although there are factor of ~2 differences in each of the resulting fitting parameters (see Table II), both sets of results are consistent with the general trends discussed in the text regarding the comparison of the fitting parameters from all three sediments.

1999). Interestingly, however, when such profiles are compared with results from shallow sediments (Fig. 8), one observes a general similarity in the exponentiallike shape of the profiles, in spite of significant differences in both the depth and concentration scales for the profiles. Examining these observations in the context of the proportionality between [DOCJooand Cox in Eq. [3] leads to the conclusion that the increasing depth scale over which organic matter remineralization occurs in these sediments (e.g., z*) likely plays a major role in explaining these observations. This point is further reinforced by the fact that at least between site M3 in Chesapeake Bay and site C on the North Carolina continental slope Cox values decrease (~20 vs ~5 mmol m~^ day~^) as [DOC]oo values increase (Fig. 8); Burdige and Zheng, 1998; Alperin et al, 1999). The results from these cores can be used to further examine suggestions posed in the previous section regarding the controls on pore-water DOC concentrations with depth. To do this I initially attempted to "fit" the data from these different

628

David J. Burdige

cores to Eq. [A-2]. However, taking this approach led to the observation that the fit of these data to this equation was insensitive to several of the model fitting parameters (e.g., R^, A, and A:H). In part this occurs because all but one of the exponential terms in Eq. [A-2] decay rapidly with depth, and therefore their actual values have a minimal impact on the overall fit of the data to this equation (see Burdige and Martens, 1990, for a discussion of a similar problem encountered in fitting dissolved free amino acid pore water profiles to an analogous equation). I am currently working to more tightly constrain the parameters in the ANS model with other measured quantities that define sediment carbon remineralization, and therefore limit the number of adjustable, fitting parameters in the model (e.g., see similar discussions in Alperin et al, 1999). Given these observations, Eq. [A-2] can be rewritten by simply removing these terms, yielding [DOC] = (Co - 25)^"^" + 25,

[4]

and since the parameters a and R^o cannot be separated from one another in the equation for Qs (Eq. [A-8]), Eq. [4] actually has only two fitting parameters (A:p [in Eq. [A-9] for y], and the combined parameter aRoo), assuming Co can be fixed with the bottom water value. Equation [4] is very similar to an analogous equation derived by Alperin et al. (1999), although here by explicitly defining the kinetics of sediment DOC production and consumption I am also able to estimate the rate constant for total DOC consumption, which is also essentially the rate constant for the consumption of pLMW DOC. The results of fitting the data from these three sites to Eq. [4] are shown in Fig. 8, where it can be seen that this modified equation does a reasonably good job of fitting data from these very different marine sediments. The resulting rate parameters from these fitting efforts are listed in Table II. As a first observation, I note that the k^ value for Chesapeake Bay sediments is ^ 2 to more than 3000 times smaller than analogous first-order rate constants for the decomposition of monomeric dissolved organic compounds such as acetate or individual amino acids (see Henrichs, 1993, for a summary of these results). This observation is consistent with previous discussions that the bulk of the porewater DOC pool (^pLMW DOC) represents relatively refractory material that turns over much more slowly than components of the mLMW DOC pool. At the same time, the Chesapeake Bay rate constant is of the same order of magnitude as those determined in degradation studies in Alaska coastal sediments of synthetic glutamic acid and alanine melanoidins (0.2-0.7 and 34 (Anderson et al, 1994; Opsahl et al, 1999) are spread around the mixing line, showing that consumption and production of DOC balance. At saHnities around 33, the Opsahl et al. (1999) data are below the mixing line, while the Anderson et al. (1994) data are above the mixing line. The latter is likely caused by biological activity as shown by Wheeler et al. (1997). Their data from the Arctic Ocean section 1994 show a similar trend (see Fig. 5 in Wheeler et al, 1997), but with fewer data above the mixing line.

LeifG. Anderson

678 160

Salinity Figure 5 DOC versus salinity for the samples from the central Arctic Ocean with 5 100 years), with a limited flux of particulate organic matter from above,

DOC in the Arctic Ocean

679

resulting in a DOC decomposition rate that could be larger than the production rate by decay of particulate organic matter. The lower DOC concentrations in the Arctic Ocean deep waters do not exclude an export of terrigenous DOM to the deep waters, as this is a function of sources and sinks. However, both the low concentration of lignin oxidation products and the predominance of a marine 8^^C signature indicate that terrigenous DOM is a minor contribution to the DOC of the deep waters of the Arctic Ocean (Opsahl et al, 1999).

IV. SUMMARY OF SOURCES AND SINKS A budget of the fluxes to and from the Arctic Ocean is given in Table III, based on measured concentrations of DOC and reported volume fluxes of the different waters. This budget does not distinguish between terrigenous and marine DOC. Generally the terrigenous DOC is high in the surface waters and low in the deep waters (Opsahl et al, 1999; Fitznar, 1999). It should be noted that the DOC budget of Table III is around 15% lower than that reported by Anderson et al (1998), a result of much new high-quality data being collected during the past few years, as referred to in Section III. The uncertainties given in Table III are based on the variability in reported DOC concentrations for the different water masses. No considerations of uncertainties in volume fluxes are included. Fortunately, the largest uncertainties are in the Atlantic and deep-water volumefluxesand these waters have a fairly constant DOC concentration. Consequently, an error in the volume influx has to be compensated by a comparable error in the volume outflux and hence have a small impact on the net DOC flux out of the Arctic Ocean. Adding the in- and outfluxes of Table III gives —5 ± 9x 10^^ g C year~\ indicating that the Arctic Ocean is neither a sink nor a source of DOC considering the uncertainty in the estimate. The in situ production of marine DOC within the central Arctic Ocean has been estimated to 6.1 g C m"-^ year~^ and the in situ respiration to 8.8 g C m~^ year~^ (Wheeler et al, 1997). Combining these numbers with the area of the deep central Arctic Ocean (5.8 x 10^^ m^) gives a total m^to production of 35 x lO^^gCyear"^ and a total in situ respiration of 51 x 10^^ g C year~^ These numbers are based on one summer investigation in a limited area and the uncertainties must be significant when applying them to a whole year and the whole central Arctic Ocean. Nevertheless, it is interesting to note that the in situ respiration of DOC exceeds that of in situ production of marine DOC, while the latter is of the same order as the added terrigenous DOC (35 x 10^^ g C year~^ relative to 23 x 10^^ g C year~^). These results indicate that the in situ respiration of DOC in the central Arctic Ocean will quantitatively consume all marine DOC produced in the central Arctic Ocean and some of that added by river runoff. This estimate does not include the DOC produced by the biota on the shelves.

680

LeifG. Anderson Table III A Budget of the DOC Fluxes to and from the Arctic Ocean Water mass

Volume flux (Sv)

DOC (AtM)

Atlantic water Deep water Pacific water Runoff

2.5 0.58 0.83 0.11

58 ±5^^ 53 ± 5 ^ 71 ± 20^ 555 ± 50^

Total in

4.02

Organic carbon transport, (1012• g C year-i)

In

Out Sea ice From EB:

From CB:

Total out Net outflow

- Surface mixed layer - Halocline - Atlantic layer - Deep water - Surface mixed layer - Halocline - Atlantic layer - Deep water

0.11 0.165 0.25 0.9 0.42 0.362 0.54 0.698 0.575 4.02

55 12 22 23

±5 ±1 ±6 ±2

112 ± 8 316 ±50^ 82 ± 15/ 70 ± 6 / 58 ± 4 / 51 ± 5 ^ 100 ± 10/

75 ± nf 53d=4/ 55 lb 5^

13 ± 2 5±1 1± 1 20 ± 1 8±1 14 ± 1 15 ± 2 14 ± 1 12 ± 1 107 ± 4 -5

±9

Note. The volume fluxes of the water masses are from Anderson et al (1998), while the DOC concentrations are means of literature values. The river runoff includes all continental freshwater input, and outflows are from the Eurasian Basin (EB) and Canadian Basin (CB), respectively. As discussed in the text, errors in the organic carbon transport figures do not include errors in the volume fluxes. ^Mean of Wheeler et al (1997), Opsahl et al (1999), and (Bussmann and Kattner, 2000). ^Data in the Greenland Sea at 1800 m (Opsahl et al, 1999). ^Mean of Walsh et al (1997), Wheeler et al (1997), and Guay et al (1999). ^The mean of the regression lines at S = 0 of Figures 2B-D. ^Melnikov (1997). /Wheeler effl/. (1997). ^Mean of Opsahl et al (1999) and Bussmann and Kattner (2000).

Even if the Arctic Ocean itself is neither a sink nor a source of DOC there is a significant export of DOC to the North Atlantic. This flux (29 x 10^^ g C year"^) is a combined result of inflow from the Pacific Ocean (22 x 10^^ g C year"^), river runoff (23 x 10^^ g C year~^), and the difference between in situ production (35 X 10^^ g C year"^) and respiration (51 x 10^^ g C year"^) within the Arctic Ocean. The above arguments together with the balanced budget support the idea that terrigenous DOC is relatively stable within the Arctic Ocean.

DOC in the Arctic Ocean

681

The finding that the outflowing deep water DOC concentrations are lower than (or very similar to) the inflowing water concentrations indicate (/) that little terrigenous DOM is exported to deep layers (as was also concluded by Opsahl et al, (1999) on the basis of lignin analysis) and (//) that little net export of marine DOM occurs to deep layers. The latter statement is supported by the arguments above that most marine DOC produced in the central Arctic Ocean is respired in the surface layers. The fact that little terrigenous DOC is exported to the deep waters of the Arctic Ocean through dense plumes originating on the shelves, where they are initiated by brine drainage from sea ice production, is an important finding as it put constraints on the global DOC budget. With regard to the total carbon budget in and out of the Arctic Ocean, the DOC fluxes are about 5% of the total carbon fluxes, calculated as the sum of inorganic and organic carbon. However, while the dissolved inorganic carbon concentration largely has a positive correlation with salinity, the DOC concentration has a negative one. Consequently, in situ production and respiration of DOC plays a relatively more important role for the carbon cycle in the low-salinity surface waters, relative to deeper layers, and it is the surface water that is in contact with the atmosphere linking the marine carbon cycle to climate.

ACKNOWLEDGMENTS Gerhard Kattner, Annelie Skoog, and Robert Benner gave valuable comments that helped improve the present manuscript. Financial support from the Swedish Research Council is greatly acknowledged.

REFERENCES Aagaard, K., and Carmack, E. C. (1989). The role of sea ice and other fresh water in the Arctic circulation. /. Geophys. Res. 94,14,485-14,498. Anderson, L. G., Bjork, G., Holby, O., Kattner, G., Koltermann, R K., Jones, E. P., Liljeblad, B., Lindegren, R., Rudels, B., and Swift, J. H. (1994). Water masses and circulation in the Eurasian Basin: Results from the Oden 91 North Pole Expedition. J. Geophys. Res. 99, 3273-3283. Anderson, L. G., Jones, E. P., and Rudels, B. (1999). Ventilation of the Arctic Ocean estimated by a plume entrainment model constrained by CFCs. /. Geophys. Res. 104,13,423-13,429. Anderson, L. G., Jones, E. P., and Swift, J. H. (2000). Export production in the central Arctic Ocean as evaluated from phosphate deficit. Submitted for publication. Anderson, L. G., Olsson, K., and Chierici, M. (1998). A carbon budget for the Arctic Ocean. Global Biogeochem. Cycles. 12,455^65. B0rsheim, K. Y., and Myklestad, S. M. (1997). Dynamics of DOC in the Norwegian Sea inferred from monthly profiles collected during 3 years at 66°N, 2°E. Deep-Sea Res. 44, 593-601. Bussmann, I., and Kattner, G. (2000). Distribution of dissolved organic carbon in the central Arctic Ocean: The influence of physical and biological properties. /. Mar. Sys. 27,209-219.

682

Leif G. Anderson

Cauwet, G., and Sidorov, I. (1996). The biogeochemistry of Lena River: Organic carbon and nutrients distribution. Mar. Chem. 53, 211-227. Degens, E. T., Kempe, S., and Richey, J. E. (1991). Summary: Biogeochemistry of major world rivers. In "Biogeochemistry of Major World Rivers" (E. T. Degens, S. Kempe, and J. E Richey, Ed.), pp. 323-347. Wiley, New York. Fitznar, H. R (1999). D-Amino acids as tracers for biogeochemical processes in the river-shelf-oceansystem of the Arctic. Ber. Polarforsch. 334, [in German]. Fransson, A., Chierici, M., Anderson, L. G., Bussman, I., Kattner, G., Jones, E. P., and Swift, J. H. (2001). The importance of shelf processes for the modification of chemical constituents in the waters of the eastern Arctic Ocean. Com. Shelf Res. 21, 225-242. Gosselin, M., Levasseur, M., Wheeler, P. A., Homer, R. A., and Booth, B. C. (1997). New measurements of phytoplankton and ice algal production in the Arctic Ocean, Deep-Sea Res. II44, 1623-1644. Gordeev, V. V., Martin, J. M., Sidorov, I. S., and Sidorova, M. V. (1996). A reassessment of the Eurasian river input of water, sediment, major elements, and nutrients to the Arctic Ocean. /. Am. Sci. 296, 664-691. Guay, C. K., Klinghammer, G. P, Falkner, K. K., Benner, R., Coble, P G., Whitledge, T. E., Black, B., Bussel, F. J., and Wagner, T. A. (1999). High-resolution measurements of dissolved organic carbon in the Arctic Ocean by in situ fiber-optic spectrometer. Geophys. Res. Lett. 26,1007-1010. Hansen, D. A., and Carlson, C. A. (1998). Net community production of dissolved organic carbon. Global Biogeochem. Cycles. 12,443^53. Hansen, D. A., Whitledge, T. E., and Goering, J. J. (1993). Patterns of nitrate utilization and new production over the Bering-Chukchi shelf. Cont. Shelf Res. 13,601-628. Hulth, S., Hall, P O. J., Blackburn, T. H., and Landen, A. (1996). Arctic sediments (Svalbard): Pore water and solid phase distributions of C, N, P and Si. Polar Biol. 16,447^62. Hulthe, G., and Hall, P. (1997). Benthic carbon fluxes—DOC versus ECO2 in shelf, slope and deep-sea environments, and relation to oxygen fluxes. Rep. Polar Res. 22^, 115-116. Jones, E. P., Anderson, L. G., and Swift, J. H. (1998). Distribution of Atlantic and Pacific waters in the upper Arctic Ocean: Implications for circulation. Geophys. Res. Lett. 25,765-768. Jones, E. P., Rudels, B., and Anderson, L. G. (1995). Deep waters of the Arctic Ocean: Origin and circulation. Deep-Sea Res. 42,131-160. Kattner, G., Lobbes, J. M., Fitznar, H. P, Engbrodt, R., Nothig, E.-M., and Lara, R. J. (1999). Tracing dissolved organic substances and nutrients from the Lena River through Laptev Sea (Arctic). Mar Chem. 65, 25-39. Lara, R. J., Rachold, V., Kattner, G., Hubberten, H. W, Guggenberger, G., Skoog, A., and Thomas, D. N. (1998). Dissolved organic matter and nutrients in the Lena River, Siberian Arctic: Characteristics and distribution. Mar Chem. 59, 301-309. Lobbes, J. M., Fitznar, H. P., and Kattner, G. (2000). Biogeochemical characteristics of dissolved and particulate organic matter in Russian rivers entering the Arctic Ocean. Geochim. Cosmochim. Acta. 64, 2973-2983. Macdonald, R. W, Carmack, E. C , McLaughlin, F. A., Iseki, K., Macdonald, D. M., and O'Brien, M. C. (1989). Composition and modification of water masses in the Mackenzie shelf estuary. J. Geophys. Res. 94,18,057-18,070. Melnikov, I. A. (1997). "The Arctic Ice Ecosystem." Gordon and Breach Science Pubhsher, The Netherlands. Menard, H. W., and Smith, S. M. (1966). Hypsometry of ocean basin provinces. /. Geophys. Res. 71, 4305^325. Olsson, K., and Anderson, L. G. (1997). Input and biogeochemical transformation of dissolved carbon in the Siberian shelf seas. Cont. Shelf Res. 17, 819-833. Opsahl, S., and Benner, R. (1997). Distribution and cycling of terrigenous dissolved organic matter in the ocean. Nature. 386,480-482.

DOC in the Arctic Ocean

683

Opsahl, S., and Benner, R. (1998). Photochemical reactivity of dissolved lignin in river and ocean waters. Limnol. Oceanogr. 43,1297-1304. Opsahl, S., Benner, R., and Amon, R. M. W. (1999). Major flux of terrigenous dissolved organic matter through the Arctic Ocean. Limnol Oceanogr. 44,2017-2023. Pocklington, R., (1987). "Arctic rivers and their discharge," Vol. 64, pp 261-268. Mitt. Geol.-Palaontol. Inst. Univ., Hamburg. Rudels, B., Anderson, L. G., and Jones, E. P. (1996). Formation and evolution of the surface mixed layer and halocline of the Arctic Ocean. /. Geophys. Res. 101, 8807-8821. Rudels, B., Jones, E. P., Anderson, L. G., and Kattner (1994). On the intermediate depth waters of the Arctic Ocean. In "The Polar Oceans and Their Role in Shaping the Global Environment" (O. M. Johannessen, R. D. Muench, and J. E. Overland, Eds.), pp. 3 3 ^ 6 . American Geophysical Union, Washington, DC. Sakshaug, E., and Skjoldal, H. R. (1989). Life at the ice edge. Ambio. 18,60-67. Schauer, U., Muench, R., Rudels, B., and Timokhov, L. (1997). The impact of eastern Arctic shelf waters on the Nansen Basin intermediate layers. J. Geophys. Res. 102, 3371-3382. Schlosser, P., Bauch, D., Fairbanks, R., and Bonisch, G. (1994). Arctic river-runoff: mean residence time on the shelves and in the halocline. Deep-Sea Res. 41,1053-1068. Slagstad, D., and Wassmann, P. (1996). Climate change and carbon flux in the Barents Sea: 3-D simulations of ice-distribution, primary production and vertical export of particulate organic carbon. Mem. Nat. Inst. Polar Res. 51,119-141. Smith, R. E. H., Gosselin, M., Kudoh, S., Robineau, B., and Taguchi, S. (1997). DOC and its relationship to algae in bottom ice conununities. /. Mar. Syst. 11,71-80. Swift, J. H., Takahashi, T, and Livingstone, H. D. (1983). The contribution of the Greenland and Barents Seas to the deep water of the Arctic Ocean. /. Geophys. Res. 88, 5981-5986. Telang, S. A., Pocklington, R., Naidu, A. S., Romankevich, E. A., Gitelson, L I., and Gladyshev, M. L (1991). Carbon and Mineral Transport in Major North American, Russian Arctic, and Siberian Rivers: The St. Lawrence, the Mackenzie, the Yukon, the Arctic Alaskan Rivers, the Arctic Basin Rivers in the Soviet Union, and the Yenisey. In "Biogeochemistry of Major World Rivers" (E. T. Degens, S. Kempe, and J. E. Richey, Eds.), pp. 75-104. Wiley, New York. Thomas, D. N., Lara, R. J., Eicken, H., Kattner, G., and Skoog, A. (1995). Dissolved organic matter in Arctic multi-year sea ice during winter: Major components and relationship to ice characteristics. Polar Biol. 15,417-4S3. Thurman, E. M. (1985). Aquatic humic substances. In "Organic Geochemistry of Natural Waters," pp. 273-361. Nijhoff/Junk PubHshers, Dordrecht. Walsh, J. J., Dieterle, D. A., MuUer-Karger, F. E., Aagaard, K., Roach, A. T, Whitledge, T. E., and Stockwell, D. (1997). CO2 cycling in the coastal ocean. II. Seasonal organic loading of the Arctic Ocean from source waters in the Bering Sea. Cont. Shelf Res. 17,1-36. Wheeler, P. A., Watkins, J. M., and Hansing, R. L. (1997). Nutrients, organic carbon and organic nitrogen in the upper water column of the Arctic Ocean: Implications for the sources of dissolved organic carbon. Deep-Sea Res. II44,1571-1592.

This Page Intentionally Left Blank

Chapter 15

DOC in the Global Ocean Carbon Cycle Dennis A. Hansell^ Bermuda Biological Station for Research, Inc., St. Georges Bermuda

I. Introduction II. Distribution of DOC A. Spatial Variability at the Basin Scale B. Temporal Variability III. Net Community Production of DOC A. Evidence for Net Production of DOC B. Regional and Global Estimates for Net Production of DOC

C. Nutrient Depletion and Net Production of DOC IV. Contribution of DOC to the Biological pump A. Evidence for DOC Export B. Exportable DOC V. Research priorities VI. Summary References

L INTRODUCTION Dissolved organic carbon (DOC) makes up the second largest of the bioreactive pools of carbon in the ocean (680-700 Pg C; WilHams and Druffel, 1987; Hansell and Carlson, 1998a), second to the very large pool of dissolved inorganic carbon (38,000 Pg C). The size of the reservoir, as well as its positions as a sink for autotrophically fixed carbon and as a source of substrate to microbial heterotrophs, ^ Present address: University of Miami, Rosenstiel School of Marine and Atmospheric Science, 4600 Rickenbacker Causeway, Miami, FL 33149. Biogeochemistiy of Marine Dissolved Organic Matter Copyright 2002, Elsevier Science (USA). All rights reserved.

685

686

Dennis A. Hansell

indicates that DOC plays a central role in the ocean carbon cycle. But what is this role, how is it realized, and what are its mechanisms and controls. The fundamentals of these questions have remained unchanged over the past 40 years (Duursma, 1962) and continue to challenge the ocean carbon research community today. A considerable amount of financial and intellectual capital has been expended and significant progress has been made over the past decade. DOC concentrations in the ocean range from a deep-ocean low of 34 JJLM to surface-ocean highs of >90 jxM (Section II). Biological processes set up this vertical gradient (net production at the surface and net consumption at depth), while certain physical conditions (high vertical stability) are required to maintain the gradient (Section II.A). The bulk DOC in the ocean can be resolved into at least three fractions, each qualitatively characterized by its biological lability (see Carlson, Chapter 4). All ocean depths contain (1) the very old, biologically recalcitrant DOC (see Bauer, Chapter 8). Its distribution is thought to be fairly uniform in the ocean, largely comprising the DOC of the deep ocean. Built upon the recalcitrant DOC at intermediate and upper layer depths is (2) material of intermediate (or semi-) lability (months to years). It is this material that is produced in the surface ocean and then mixed into the main thermocline, thereby reducing the vertical concentration gradient and contributing to carbon export (Section IV). Concentrations of this fraction can be 10-30 /xM in the upper ocean, and near zero in the deep ocean. The surface ocean alone contains (3) the highly biologically labile fraction of DOC, with lifetimes of days to months and concentrations of just a few to tens of micromolar of C. This latter material is most important for supporting the microbial heterotrophic processes in the ocean (see Carlson, Chapter 4) and shows high variability seasonally. In this chapter, the role of DOC in the ocean carbon cycle is considered in its broadest temporal and spatial scales. The chapter begins with an evaluation of the spatial distribution of DOC at the regional and basin scales, in both the surface and deep ocean. In this context, the distribution of DOC relative to the distribution and timing of marine productivity is evaluated. The older data sets reporting DOC distributions are appraised here as well. The next Section evaluates temporal variability, with consideration of how DOC varies seasonally from high to low latitudes. Following the assessment of variability, the net community production of DOC is examined. The focus is on DOC that accumulates for durations with biogeochemical relevance. This Section is followed by an evaluation of the contribution of DOC to the biological pump. We examine the mechanisms and locations of DOC export, and thus develop an understanding of the controls on export. The chapter concludes with priorities for present and future research, as well as a brief synthesis of the findings reported. Note that organic carbon in the ocean is distributed between the dissolved and particulate (POC) fractions. Summed, these fractions are referred to as total organic carbon (TOC). It is common to measure TOC directly in the water column

DOC in the Global Ocean Carbon Cycle (analysis of unfiltered water), even when DOC is the pool of interest, when the POC concentrations are very low relative to the DOC concentrations. This situation is common at ocean depths well below the surface layer (at depths >200 m), as well as in some surface ocean regions where POC concentrations are normally a few micromolar. The latter conditions are found in the oligotrophic ocean and in highlatitude systems during winter. In these situations (deep water and low-POC surface water), TOC serves as a very close approximation to the DOC concentrations. A primary reason for measuring TOC in these waters, rather than measuring DOC directly, is to avoid contamination by handling (filtering, transfers, etc.) the sample. The term DOC is used in this chapter both for true DOC analyses, and when TOC was measured in deep or POC-impoverished surface waters. The term TOC is reserved for use when DOC and POC, measured separately, are sunmied.

11. DISTRIBUTION OF DOC A. SPATIAL VARIABILITY AT THE BASIN SCALE During the decades leading up to the 1990s, DOC data were relatively sparse because relatively few laboratories made the measurements. An early body of work that stands as providing some of the greatest spatial coverage of an ocean region is that by Duursma (1962). He conducted extensive DOC surveys in the northern reaches of the Gulf Stream and its offshoots south of Greenland, finding the spatial variability associated with hydrographic features that we would likely find today. Since his early work, the few additional ocean sections occupied in the decades leading to the 1990s produced data of uncertain quality (see discussion by Wangersky, 1978, and below). Consequently, our sense for the distribution of DOC in the ocean has been highly uncertain. In this discussion, findings from recently occupied sections in various ocean basins will be discussed (locations identified in Fig. 1) and some of the older sections evaluated. 1. Upper Ocean Distributions Meridional sections from the eastern and western South Pacific and the central Indian Ocean show that the highest upper ocean DOC concentrations are typically found in the low to mid latitudes (Fig. 2 [see color plate]). Concentrations decrease into colder water, whether as a horizontal gradient along the surface from low to high latitudes or vertically with increasing depth. Vertical stability provided by the main thermocline of the open ocean supports the accumulation of DOC in the surface waters. Where vertical stability is strong, DOC concentrations are relatively high; where stability is weak, DOC concentrations can remain at low levels (Fig. 2c). The cold, deep waters have the lowest concentrations, and where

687

688

Dennis A. Hansell

Figure 1 Map depicting locations in the global ocean from which data are shown in this chapter. The solid Unes (gray) represent ocean sections; the triangles are the sites of the time-series stations near Bermuda and Hawaii; the filled circles and the triangles are the sites of the deep-water DOC analyses in addition to the time series sites.

these waters ventilate at high latitudes, similarly low DOC values are present (Figs. 2b and 2c). Where low-DOC, subsurface water mixes to the surface its impact is felt in the surface DOC concentrations. Upwelling sites, both in coastal regions and along the equator, normally have relatively low DOC values at the surface where upwelling is strongest. In the central Equatorial Pacific, DOC is depressed at the surface because of upwelling (note the upward doming of the subsurface DOC contours at the equator; Figs. 2b and 2c). In the central Indian Ocean, where equatorial upwelling is weak, DOC is rather uniform from the subtropical gyre to across the equator (Fig. 2a). DOC along the equator in the Pacific shows the controls by hydrography and biology (Fig. 3 [see color plate]). The Equatorial Undercurrent, near 200 m west of the dateline, has a DOC concentration of ~55 /xM (Hansell et al, 1997b). This water is transported to the east, shoaUng to near surface in the central and eastern Equatorial Pacific, bringing with it low-DOC water. The return flow of surface water to the west undergoes an increase in DOC (to ~65 /xM) due to biological activity. The highest DOC concentrations in Fig. 3 (>70 /xM C; largely west of 165° W in the surface 100-120 m) are associated with the Western

DOC in the Global Ocean Carbon Cycle Pacific Warm Pool (Hansell et al, 1997b; Hansell and Feely, 2000). The front separating the DOC-enriched Warm Pool to the west and the recently upwelled water to the east varies with the ENSO state, being found further to the west during La Nina conditions (Dunne et al, 2000). The impact of upwelling on equatorial DOC concentrations exists at coastal upwelling sites as well. Along the coast of Oman in the Arabian Sea, strong upwelling occurs during the Southwest Monsoon. Low surface DOC concentrations are present in coastal water during such periods although productivity can be quite high (Hansell and Peltzer, 1998). UpwelUng along the northwest margin of the African continent similarly forces a shoaling of the DOC isolines (Postma and Rommets, 1979). Similarly, Doval et al (1997) reported a decrease in subsurface DOC in northwest Spain due to upwelling. Ocean margins influenced more by riverine inputs than by upwelling tend to show increases in DOC concentrations. Rivers introduce water with high DOC concentrations (see Cauwet, Chapter 12), thus raising concentrations along the coast. One example is in the Chesapeake Bay outfall, where DOC concentrations increase from 70 ^xM in the surface Sargasso Sea to >200 JJM in the Chesapeake Bay mouth (Bates and Hansell, 1999). Guo et al (1994) reported onshore DOC concentrations of 131 /xM off Galveston, Texas, and moderate concentrations of 83 /xM offshore in the Gulf of Mexico. Property/property plots of DOC and salinity show the conservative nature of riverine DOC as it mixes with oceanic water. In general, the strength and direction of concentration gradients between the surface open ocean and the coastal ocean depend on the degree of upwelling of low-DOC water from below or invasion of DOC-enriched freshwater from the continent. Comparing two zonal sections in the North Atiantic provides further evidence for the control physical properties of the water column play on DOC distributions (Fig. 4 [see color plate]). A Section at 24°N shows strongly enhanced DOC concentrations in the upper 200 m (up to 80 /xM C), reflecting the strong stratification present in the subtropical gyre. In a more northerly Section, surface DOC is lower (>60 jjM C) and the concentration contours are pushed deeper into the water column. Note, for example, the 55 /xM DOC contour at 200-300 m along 24°N, but at 200-600 m on the northern Section. This change in contour depths reflects the weaker stratification at higher latitudes, and subsequent downward mixing of the surface produced DOC. 2. Deep-Ocean Distributions Reports on the distribution and variability of DOC in the deep ocean have been conflicting. Measurements from the 1960s (discussed below) suggested strong, horizontal gradients in DOC. More recently, Druffel et al (1992) reported a modest 5 /xM concentration difference between the deep waters near Bermuda and Hawaii. Martin and Fitzwater (1992), in contrast, reported the complete absence of DOC

689

690

Dennis A. Hansell

gradients in the deep ocean. Hansell and Carlson (1998a), in an effort to narrow the uncertainty, surveyed representative sites in the deep ocean (Fig. 1). They found a 29% decrease in DOC concentration from the northern North Atlantic (48 /xM in the Greenland Sea) to the northern North Pacific (34 /JM in the Gulf of Alaska) (Fig. 5, top). The gradient reflects the export of DOC-enriched (formerly subtropical) water during North Atlantic deep water (NADW) formation (Fig. 5, bottom) and the decrease in DOC (by mineralization and mixing) along the path of deep ocean circulation away from the North Atlantic. The formation of Antarctic bottom water (AABW) does not introduce additional DOC to the deep ocean (see Section IV), so the concentrations remain low near those sites. The small increase in DOC concentrations from the Southern Ocean into the deep South Pacific and Indian oceans is enigmatic and the source unidentified (Hansell and Carlson, 1998a). Possible causes include inputs from marginal seas (Red Sea, Arabian Sea, and Bay of Bengal for the Indian Ocean), inputs due to dissolution of sinking biogenic particles, non-steady-state conditions in the deep-ocean concentration gradients of DOC, and, of course, unidentified processes. The highest deep-water DOC concentrations may be those in the deep Eurasian Basin of the Arctic Ocean (see Anderson, Chapter 14), where concentrations >50 /j^M C have been reported (Anderson et aL, 1994). The sources of this material must be terrestrial runoff and Arctic continental shelf produced DOC (Opsahl et aL, 1999; Wheeler etai, 1997). It is interesting to speculate as to the mechanisms responsible for DOC concentration decreases in the deep ocean. Certainly microbial mineralization and mixing contribute, but, based on our present knowledge, these mechanisms appear to be inadequate. The DOC concentration decreases by 14 /xM over the length of the deep limb of the "global conveyor belt," but would the marine microbes we are most familiar with today be satisified with such meager rations over the half millennium required for transport over that distance? The apparent rate of oxidation (^30 nM year"^ over ^500 years), and the amount of energy derived over these several centuries, is miniscule. Perhaps the poorly understood Archaea, now known to inhabit the deep ocean, are designed to catabolize recalcitrant DOC at such low rates. Perhaps microbes play only a secondary role, and DOC is removed primarily by coagulation and formation of sinking particles, or it is stripped from the water column by particles passing through the water colunm. The true mechanisms for DOC loss need to be resolved. 3. Relation to Productivity Given the surface DOC distributions described here (Fig. 5), of low DOC near sites of upwelling or deep mixing and high values in stratified water, a general observation can be made: upper ocean DOC concentrations are relatively high in oligotrophic waters where regenerated production dominates, and low in systems

691

DOC in the Global Ocean Carbon Cycle 50

Pacific/Indian Sector

Indian

32' 80°N

40"N

80°S

40"S

40°S

40°N

80"N

Latitude

Atlantic Sector

Pacific/Indian Sectors

av^ace NPOH^//

AABW

North

Circumpolar water

South

CDW North

Figure 5 (Top) Distribution of DOC in the deep-ocean. The x-axis is viewed in the context of the deepocean circulation, with fonnation in the North Atlantic, circulation around the Southern Ocean, and flow northward into the Indian and Pacific oceans. Station locations in Fig. 1. (Bottom) The general patterns of ocean circulation driving the deep ocean DOC signal. DOC-enriched surface water is introduced to the deep ocean in the North Atlantic. This water moves south as North Atlantic deep water (NADW), to the circumpolar waters of the Southern Hemisphere. DOC-impoverished Circumpolar deep water (CDW) flows north into the Pacific and Indian oceans. Deep return flow to the North Atlantic is via Antarctic bottom water (AABW) and to Antarctica via North Pacific (NPDW) and Indian Ocean deep waters (lODW).

692

Dennis A. Hansell

where new nutrients are introduced to the surface. Such a meridional gradient has been reported for the Equatorial Pacific (Tanoue, 1993), the South Pacific (Hansell and Waterhouse, 1997; Doval and Hansell, 2000) and the North Atlantic (Duursma, 1962; Kahler and Koeve, 2001) and is evident in Fig. 2. The regenerated vs new production nature of these systems is a reflection of the conmiunity compositions within them. The mechanisms by which conununity composition controls DOC concentrations are not understood (see Carlson, Chapter 4). DOC concentrations are also controlled by the vertical stability of the water colunm. The highest DOC concentrations in the open ocean are normally found where stratification of the water column is highest (Hansell and Waterhouse, 1997; Hansell and Feely, 2000). This finding suggests that stability facilitates the retention of DOC in the upper ocean. The lowest DOC concentrations, to the contrary, exist where DOC-depleted subsurface water is introduced to the surface, either by vertical mixing or upwelling. These high nutrient sites can experience large but brief seasonal increases in DOC concentrations, however (see below). Because of the role of ocean stratification in controlling DOC concentrations, a positive correlation between DOC concentrations and primary productivity (an oft predicted relationship) is absent in much of the oligotrophic, open ocean. Menzel and Ryther (1970) reported the absence of this correlation early and evidence for the generality will be given using data from the Sargasso Sea later in the chapter (Section II.B.2). In fact, in the highly stratified portions of the open ocean, DOC broadly correlates positively with temperature (Hansell and Waterhouse, 1997; Doval and Hansell, 2000), another sign of the importance of physical control on concentrations. At higher latitudes, however, where DOC concentrations are depressed during the winter, elevated DOC values indeed follow springtime elevation of primary productivity (Borsheim and Myklestad, 1997; Chen et al, 1996; Carlson et al, 2000). This positive relationship between primary production and DOC was reported early by Duursma (1963) and has been discussed elsewhere (Williams, 1995). In high-latitude systems, increased water colunm stability favors both phytoplankton growth and DOC accumulation in the upper ocean. The data indicate that low-latitude, highly stratified environments behave very differently than high-latitude environments in terms of the coupling between DOC dynamics and primary production. So, while their observations are in apparent conflict, both Menzel and Ryther (1970) and Duursma (1963) were correct about the relationship between DOC and productivity; but they were correct specifically for the hydrographic systems they were evaluating. 4. Historical Data With the onset of discussions surrounding the use of the high-temperature combustion (HTC) systems for DOC analysis (Sugimura and Suzuki, 1988), much attention has been paid to whether or not the earlier data are accurate and, therefore,

DOC in the Global Ocean Carbon Cycle of value (Sharp, 1997). A comparison of what we find in the ocean today with that reported in earlier decades shows some older data and findings to have serious flaws. A comparison of historical and recent data from the surface ocean cannot be easily made because of the wide natural variability in those waters (see below). The most useful comparisons between historical and recent data are made in the intermediate and deep ocean, where significant changes in concentration over a few decades (the sampling interval) are unlikely. Menzel (1964) reported DOC concentrations in the intermediate depths (400800 m) of the Arabian Sea and western Indian Ocean to range from 0.4 to 1.6 mg/L (30 to 130 /xM DOC). This wide range is not reproducible anywhere in the intermediate or deep ocean using modem techniques, nor was it evident during the US Joint Global Ocean Flux (US JGOFS) program in the Arabian Sea during 1995 (Hansen and Peltzer, 1998). Menzel and Ryther (1970) also reported a very unlikely DOC concentration doubling at all depths > 1000 m between the waters northeast and southeast of South America. Romankevich and Ljutsarev (1990), reviewing investigations conducted by the Soviet Union, reported DOC off Peru at 500-1000 m to be an unlikely 1 mg/L (^83 /xM). Soviet measurements in the deep Bay of Bengal exceeded 1 mg/L as well. These latter DOC concentrations are probably high by a factor of two. Williams et al (1980) reported DOC concentrations in the central North Pacific a few meters off bottom (5650 m) that were elevated by twofold relative to the values at 2000-5000 m. Such a strong DOC gradient, indicative of sedimentary input of DOC to the bottom layer, has not been confirmed using modem techniques and extensive near-bottom surveys. Recent data, using modem HTC techniques, should be viewed with caution as well. Dileep Kumar et al. (1990) reported a strong DOC concentration gradient from the central Arabian Sea to the westem Indian Ocean, increasing from 100 to 300 /zM at 2500 m. The high DOC concentrations and wide range reported are unlikely to be accurate representations of that system.

B. TEMPORAL VARIABILITY The temporal variability of DOC concentrations in the surface ocean has been noted since the earliest days of the measurement. Duursma (1963) reported a twofold increase in DOC concentrations in the North Sea, from winter lows of 0.8 mg/L (~66 jiM) to spring and early summer highs of 1.8 mg/L (~150 /xM). The increase in DOC concentrations started some weeks after the spring phytoplankton bloom. Holmes et al. (1967) reported large spikes in DOC concentrations, from a baseline of 1 mg/L up to 4-5 mg/L (330-415 /xM), during several red water dinoflagellate blooms off La JoUa Bay, Califomia. Here, too, the DOC peaks followed the decline of the blooms. Williams (1995) evaluated the seasonal accumulation of DOC using data from Parsons et al. (1970), Banoub and WiUiams (1973),

693

694

Dennis A. Hansell

and Duursma (1963), suggesting that the accumulation of C-rich dissolved organic matter resulted from nitrogen limitation. The possible role of nutrient depletion in the generation of DOC is discussed below. (See Carlson, Chapter 4, for a more complete listing of publications reporting temporal variability of DOC.) 1. High Latitudes Strong seasonal increases in DOC concentrations associated with phytoplankton blooms appear to be characteristic of systems that receive high input of new nutrients over the winter periods. The waters of the Ross Sea polynya, for example, undergo deep mixing over the winter such that nitrate concentrations exceed 30 fiM prior to the spring bloom (Bates et ai, 1998). DOC concentrations increase in the surface layer from winter lows of 42 /xM to summer highs of 65-70 /xM (Carlson etai, 1998). High southern latitude systems can experience large increases in DOC concentrations (tens of micromolar C), with the wintertime baseline concentration as low as the much deeper waters (Carlson et aL, 2000; Wiebinga and de Baar, 1998; Kahler et aL, 1997). The Ross Sea undergoes DOC concentration increases of 15-30 /xM where the Phaeocystis and diatom blooms are particularly strong (Carlson etal, 2000). Where the blooms are small because of various controls on plant growth (deep mixing, iron limitation, etc.), the DOC concentrations remain low (e.g., over the Ross Sea shelf break, with a gain of 2x increase in DOC concentrations during the summer. It is apparent, though, that the winter lows of DOC concentrations in the high northern latitudes are not as low as the local deep-water values (in contrast to the conditions found in the Southern Ocean). This finding holds true along 20°E in the North Atlantic, where Kortzinger et aL (2001) reported the winter low DOC to be 53 fiM C, well above the deep-ocean values in the region. The more physically stratified nature of the northern systems prevents full water column overturn and homogenization of the DOC each winter. 2. Mid-latitudes The more oligotrophic, mid-latitude zones of the ocean do not show the same seasonality (in either strength or direction) as the high latitudes or other nutrientrich areas. In the Sargasso Sea, where convective overturn during the winter introduces small amounts of new nutrients to the euphotic zone and phytoplankton blooms follow (Michaels and Knap, 1996), the seasonality of DOC in the surface ocean contrasts that found at high latitudes (Carlson et aL, 1998; Hansell and Carlson, 2001a). Overturn of the water column coincides with the spring bloom

DOC in the Global Ocean Carbon Cycle there because adequate light is present at these mid latitudes. The effect is to mix low DOC subsurface water upward, thereby reducing the DOC concentrations during the periods of highest primary productivity. Once stratification reasserts itself with warming of the surface ocean, and the bloom terminates, DOC concentrations rebuild to normal summer levels (Fig. 6 [see color plate]). The concentration change from the annual low to the annual high is only 3-6 /xM, a very small range compared to high-latitude systems. The lowest winter concentrations remain well above the deep-water values. While the DOC concentrations in the Sargasso Sea are lowest during the winter overturn/spring bloom period, the same cannot be said for the integrated DOC stocks. Relatively deep convective overturn maintains the low surface DOC concentrations but the bloom still supports the net production of as much as 1-1.5 mol m"^ of DOC over the upper 250 m (Fig. 7; Carlson et al, 1994; Hansen and Carlson, 2001a [see color plate]). This increase in DOC stock is as large as that seen in the much more productive Ross Sea (Carlson et al, 2000). DOC and bloom dynamics in the Arabian Sea during the NE Monsoon are similar to that in the Sargasso Sea. Convective overturn in the Arabian Sea, forced by cool dry winds off the Tibetan plateau, mixes moderate amounts of nutrient into the euphotic zone. There, too, DOC concentration changes are not large during the bloom, but the increase in DOC stock can be 1.5-2 mol C m~^ (Hansell and Peltzer, 1998). It is interesting that while the seasonal range for DOC in the western Sargasso Sea (at ~31°N) is only 3-6 luM, the seasonal range at the same latitude in the eastern North Atlantic can reach 10-20 /xM (Kortzinger et al, 2001). The western North Atlantic is generally warmer and more stratified than in the east, suggesting differing community composition and productivity between the sites. This follows from the gyre circulation patterns: the northward flow of water in the west, from the warm equatorial region to higher latitudes, lends itself to high vertical stability and highly oligotrophic conditions; the southern flow in the east, carrying cooler water from the north, would lend itself to less stability and less oligotrophic conditions. It may be that the more stable system in the west experiences less primary productivity and net DOC production than the system in the east. Physical characteristics of the systems and the biological regimes they support are centrally important in controlling DOC dynamics. 3. Low Latitudes Low-latitude systems that do not undergo winter freshening of the surface layer do not show seasonality in DOC concentrations. The waters at Station ALOHA, north of Hawaii at 23°N, represent such a system. There, variability in DOC occurs at interannual time scales, but there is no recurring trend with seasons (Fig. 6). Church et al (2001) reported a net accumulation from 1993 to 1999 of a DOM

695

696

Dennis A. Hansell

pool that was enriched in C and N, relative to P. These long-term changes may be a manifestation of the broad, ecosystem-wide shift from N to P limitation described by Karl (1999). No such shifts have been noted in the Sargasso Sea, hence the near constancy in summer time DOC highs from year to year. Note that the surface DOC concentrations at ALOHA are much higher than the highs at BATS (Fig. 6) and higher than any values found along 24°N in the North Atlantic (Fig. 4). Why this difference exists between these similarly low latitude zones of the North Atlantic and North Pacific is unknown. Community composition may be key, but an evaluation has not been conducted. 4. Deep Ocean Whether or not there is measurable temporal variability of DOC in the deep ocean remains debatable. Hansell and Carlson (2001a) did not resolve DOC variability in the deep Sargasso Sea over 6 years of time series measurements. Similarly, Hansell and Peltzer (1998) found no variability in the deep Arabian Sea over a single year, even through periods of very high sinking particle flux. Bauer et al (1998), in contrast, reported significant (8 /xM) long-term (2-year) changes in DOC in the deep eastern North Pacific. They tied these variations to natural variability ("patchiness") and exchanges with sinking POC. Why there may be variability at this site and not at the others studied needs to be resolved. 5. Short-Term Biological Events Further variabiUty in DOC concentrations can be expected to occur with biological "events." Examples are blooms of red tide organisms described by Holmes et al (1967) and of diazotrophs. Onset of enhanced nitrogen fixation rates in openocean systems can increase DOM stocks considerably. Karl et al. (1997) reported organic nitrogen concentration increases of several micromolar which should coincide with several tens of micromolar increase in DOC. A case in point is the western tropical South Pacific, where relatively high DOC is present under the zone of the atmospheric South Pacific Convergence Zone. Hansell and Feely (2000) suggested that the excess precipitation in this system increased vertical stability, thereby favoring nitrogen fixers and in turn increasing concentrations of DON and DOC. Near the continental margins, DOC concentrations will vary with the strength of DOC-enriched riverine inputs or coastal upwelling, both of which vary seasonally (Cauwet, Chapter 12; Hansell and Peltzer, 1998). High riverine input may result in high-DOC concentrations; strong upwelling reduces the DOC concentrations. Zones of equatorial upwelling similarly exhibit the lowest DOC concentrations during strong upwelling (e.g., La Nina), and the highest values during reduced upwelling (e.g.. El Nino; Peltzer and Hayward, 1996). In this way, physical stability plays a major role in controlling DOC concentrations both along the margins, in

DOC in the Global Ocean Carbon Cycle the open ocean and in equatorial upwelling systems (Carlson and Ducklow, 1995; Hansen and Waterhouse, 1997; Tanoue, 1993). 6. Summary Our present understanding of seasonal variability in DOC can be summarized here: At high latitudes, where spring blooms are intense, we expect to see large DOC concentration changes. Because the winter DOC concentrations are low in these high-latitude systems, the highest concentrations during sunmier may be no higher than the summer highs in the low-latitude gyre systems, but the concentration change between seasons may be large. However, the large increases in concentration do not necessarily translate into large accumulations of DOC stock (vertically integrated loads of DOC) because of the normally shallow euphotic zones in these highly productive systems (high concentrations but over little depth). In mid-latitude open-ocean regions, such as the Sargasso and Arabian Seas, where convective overturn introduces moderate nutrient loads, DOC concentration ranges between seasons can be relatively small, though the change in integrated stocks can be a relatively large signal (comparable to the change in DOC stock in the Ross Sea). The overturning water column mixes the DOC too deeply for a strong surface accumulation to occur, as found in blooms occurring in more stratified systems but the small concentration change over large depths results in significant increase in stock. At mid-latitude coastal sites with significant winter recharge of surface nutrients, DOC seasonality will be strong. Upwelling reduces the DOC concentrations while high riverine inputs increase them. At low-latitude sites where spring blooms are absent, no seasonality is evident; but, as with all ocean regions, interannual variability exists. The DOC that accumulates each year at mid-latitudes has a lifetime exceeding the season in which it was produced, so it can be transported elsewhere with surface currents, or be available for export during the subsequent winter overturn events. At higher latitudes, the seasonally produced DOC is seen to have a lifetime shorter than that of the season of production; thus it undergoes net consumption by microbes once primary production is reduced with the onset of Fall conditions. This material is not as available for export (see Section IV.A).

III. NET COMMUNITY PRODUCTION OF DOC DOC is produced on a daily basis as part of the primary and secondary production systems in the surface ocean. Most of the DOC released is mineralized on the time scale of hours to days. For DOC to play a role in the ocean carbon cycle beyond serving as substrate for surface ocean microbes, it must act as a reservoir for carbon on the time scales of ocean circulation. This it does, given that the

697

698

Dennis A. Hansell

ET-^ Deep and Bottom water m High density mode water • I Low density and Subtropical mode water —— Subtropical gyre circulation

Figure 8 Distribution of sites of water column overturn (from Talley, 1999), general patterns of surface circulation in the subtropical gyres, and proposed distribution of exportable DOC. Overlap in the distribution of exportable DOC (background field of white) and sites of ocean ventilation (sites colored by gray scale) favors DOC export; a lack of overlap precludes export. The waters of the Southern Ocean (slanted stripes) are without exportable DOC present, so where these waters overlap sites of ventilation, little export is expected.

production and accumulation of DOC in the surface ocean has been demonstrated (Figs. 2-8). The rates of, and controls on, the net production of DOC, topics not well understood at this time, are the focus of this section. Because so few DOC data exist, particularly from ocean systems for which accumulation has been evaluated, it is useful to normalize estimates of DOC accumulation to a more broadly available and easily measured variable. DOC accumulation as a function of net community production (NCP) has proven useful in this way (Hansell and Carlson, 1998b; Kortzinger et aL, 2001). NCP occurs when autotrophic production exceeds heterotrophic consumption, such as during a spring bloom. It is a process that largely results in the export of carbon and new nitrogen from the euphotic zone as sinking biogenic particles and in this way is analogous to new production (Dugdale and Goering, 1967). If DOC accumulates, then DOC too is a sink for NCP. NCP is estimated most directly by measuring the biological drawdown of the reactants (dissolved inorganic carbon and/or nitrate) or as the flux of the products (i.e., accumulation of DOC, suspended POC, export of sinking biogenic particles.

DOC in the Global Ocean Carbon Cycle and contributions by migrating zooplankton). The Section above on temporal variability of DOC sheds light on the net production of DOC. As a rule, oceanic regions showing seasonality of DOC concentrations are experiencing some transfer of NCP into the DOC pool.

A. EVIDENCE FOR NET PRODUCTION OF DOC Seasonal increases of DOC stocks in the Ross Sea indicate that 8-20% of NCP in the polynya system accumulates each growing season as DOC (Bates et al, 1998; Carlson et al, 2000; Hansell and Carison, 1998b; Sweeney et aU 2000). The balance of NCP is lost to the deep ocean as sinking biogenic particles, mostly Phaeocystis and diatoms. Annual rates of NCP in the Ross Sea polynya are 6-14 mol C m"^, so net DOC production of 1.2-2 mol C m"^ occurs over the growing season (Bates et al, 1998; Carlson et al, 2000; Sweeney et al, 2000). The net production of DOC in the Ross Sea is about that in the Sargasso Sea (1-2 mol C m~^; see above), but the Sargasso Sea has a much lower annual rate of net commiunity production. The rate of DOC production in the Ross Sea, normahzed to NCP, is similar to that found in the Equatorial Pacific. Estimates of net DOC production as a percentage of NCP in the central Equatorial Pacific range from 6 to 40%, with most estimates near the 20% level (Archer et al, 1997; Hansell et al, 1997a,b; Zhang and Quay, 1997). These values from the Equatorial Pacific are similar to the Equatorial Atlantic (20%; Thomas et al, 1995), but significantly lower than prior estimates in the Equatorial Pacific by Murray et al (1994), Feely et al (1995), and Peltzer and Hay ward (1996). Those latter authors estimated net DOC production closer to 75% of NCP, but those findings have been challenged (Hansell et al, 1997b; Zhang and Quay, 1997). Noji et al (1999) suggested that more than half of NCP in the Greenland Sea accumulated as DOC, high compared to findings from other nutrient-rich sites. Alvarez-Salgado et al (2001) reported that 20% of net ecosystem production accumulated as DOC in a coastal upwelling environment along the Iberian margin in the North Atlantic. This rate is very similar to that reported for the Equatorial Pacific and the Ross Sea. Net DOC production in the Ross Sea, the Equatorial Pacific and the Iberian margin takes place when the conditions are right for net autotrophy. In the Ross Sea, this occurs when vertical stability and light are available, while in the Equatorial Pacific and the coast of Spain light becomes available following upwelling. At these three sites, vertical stability is relatively strong during the periods of net production. The Sargasso Sea contrasts those systems. Light is generally available year round but nutrients are not, so a reduction in vertical stability (convective overturn of the water column and entrainment of nutrients) is required for net autotrophy. A representative year (July 1994 to July 1995) for DOC in the Sargasso

699

700

Dennis A. Hansell

Sea is useful for demonstrating net DOC production (Fig. 7). Winter overturn and mixing of the water column was both the cause of concentration reductions and the trigger for net DOC production each year following nutrient entrainment and subsequent new production (Carlson et aL, 1994; Hansell and Carlson, 2001a). The net production of DOC at the BATS site varies interannually as a function of the maximum in the winter mixed layer depth. The greater the vertical mixing (and nutrient entrainment) in the Sargasso Sea, the greater the net production of DOC (Hansell and Carlson, 2001a). In winter 1995 (Fig. 7), the DOC stock increased by 1.4 mol C m~^in response to maximum mixing depths of 260 m (note the net production of DOC in the upper 250 m of the water colunm; Fig. 7b). In subsequent years experiencing shallower maxima in mixed layer depth (