Biogeochemistry of Marine Dissolved Organic Matter

Biogeochemistry of Marine Dissolved Organic Matter

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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

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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.

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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.

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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.

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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-^

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10-8 iiiiiii III

I

I

lo-^ mill I

III

10-^°

Dissolved Colloids

^fel Sand

I [

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_

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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

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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

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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

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200 /xL) on linearity (0-13 mM), precision (95%, with the exception of sulfathiazol (4000 amu) fractions. These results were supported by the finding that over 60% of the initial DOC (232 IJM by HTCO) in an unspecified seawater sample subjected for 2 h

17

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Figure 8 DON versus DOC in surface waters from the North Pacific and East China Sea (data from Table VII of Sugimura and Suzuki, 1988). The five open squares indicate high-DOC waters with a salinity less than 33 that were collected off the mouth of the Changjiang (Yangtze) River. The equation of the best-fit line (r^ = 0.625) to the other 32 data points (sohd squares) is DON = (0.554 ± 0.078) x D O C + (25.9 ±1.5).

to persulfate oxidation still could be measured by HTCO in the residual water. The bulk of the remaining DOC reportedly occurred in the >4000-amu size fractions, indicating again that high-molecular-weight DOC is particularly resistant to persulfate oxidation. In the same paper, Sugimura and Suzuki reported application of the new HTCO method in an extensive survey of surface seawaters and three depth profiles from the western North Pacific. In general, these field results (Fig. 8) strongly reinforced the implications of the previously described laboratory tests. DOC values in the range of 180-490 luM were measured for surface seawaters south of Japan, with values above 275 /xM being limited to five low-salinity (300 /xM DOC would have more than enough reducing potential (>380 /xM) in DOM alone to remove all the oxygen it could possibly dissolve (~350 /xM O2, at 0°C and salinity of 35; Broecker and Peng, 1982). Such waters would be highly prone to anoxia if in addition they hosted any in situ respiration of sinking organic particles. Sinking bioactive particles also would be expected to cause a pronounced downward displacement in TDN profiles, which is not evident in Fig. 9. These inconsistencies, however, were difficult to test at the time because the exact composition of the platinized alumina support (Sumitomo Chemical Industry Co. Ltd.), which appeared to be critical to the efficient performance of the Sugimura and Suzuki HTCO DOC and DON analyzers, was proprietary information. Another hurdle for comparative measurements was the authors' report that reliable DOC results could be obtained only when seawater samples were filtered and immediately analyzed

Why Dissolved Organics Matter

21

aboard ship. In addition, neither reference seawater samples nor DOM-free water for blank testing were widely available to the oceanographic community at that time. Not surprisingly, the "revolutionary" (Toggweiller, 1989a)findingsof Sugimura and Suzuki generated a rush by other research groups to acquire and apply Pt-catalyzed HTCO units for DOC and DON analyses. Not only did such instruments appear to measure DOC and DON more efficiently than predecessors employing wet chemical oxidation, they also were faster, easier, more readily automated and required much smaller sample volumes (~ 100 /xL) than were necessary (5 mL) for the standard persulfate method. By 1991, Pt-catalyzed HTCO analyzers for DOC and DON were available from at least six different companies and were in use by more than 20 research groups in the oceanographic community (Hedges et al, 1993). Data both supporting and refuting the "New" DOC/DON hypothesis flooded the oceanographic literature. Theories for why this previously unknown component of seawater DOM was measurable only by Pt-HTCO analyzers proliferated, as did papers attempting to characterize these materials and investigate their biogeochemical importance (see Hedges and Lee, 1993). Unfortunately, the results of these follow-ups were mixed, including two papers (Benner et al, 1992; Ogawa and Ogura, 1992) that cast doubt on most of Suzuki's size distribution and composition findings. Such divergent observations in the early 1990s caused confusion, consternation, and cautious support by national agencies for DOM research. The growing furor led to an NSF/NOAA/DOE-sponsored workshop on the "Measurement of Dissolved Organic Carbon and Nitrogen in Natural Waters" that was held in Seattle in July 1991 (Hedges and Lee, 1993). To assess measurement uniformity within the oceanographic conmiunity, ampoulated samples of surface, mid-depth, and deep-ocean waters collected at the Hawaii ALOHA station (and from a Hawaiian river) were distributed before the workshop to invited participants. A total of 13 independent DON measurements, and 34 independent DOC analyses of the sample suite were made. The results were not encouraging (Hedges et al., 1993). The DON measurements varied by an average of ±30% of the mean value for the samples and were not related to any known aspect of the analyzers or their use. The corresponding DOC concentrations varied by an average of ±40% of the mean values, with HTCO instruments generally measuring higher concentrations than were obtained for the same samples by more conventional wet-chemical techniques (Fig. 11). The DOC differences, however, largely disappeared when the mean value for each analyst's four individual samples was subtracted from the corresponding individual measurements (Hedges et al, 1993). Such a pattern following mean subtraction would be expected if the major sources of difference among DOC concentrations measured by the participating labs were traceable to large background signals intrinsic to each instrument and its method of operation.

22

John I. Hedges 350| 300

" Oj Minimum Zone

250| ^ 200| U Surface Ocean

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Means from Each Set of Analyses Figure 11 Trends in measured DOC among reference seawater samples analyzed for the Seattle DOC/DON Workshop (data from Hedges et ai, 1993). Individual data sets are listed in order of decreasing measured concentration for the surface sample. Analyses by wet chemical oxidation methods are indicated by shaded backgrounds.

Supporting evidence for large DOC blanks intrinsic to HTCO-based instruments came from additional workshop reports. In particular, a critical evaluation by Benner and Strom (1993) of the analytical blank associated with DOC measurements using HTCO instruments showed that the platinized-alumina packing used in the combustion columns of almost all of the tested commercially-available analyzers (Shimadzu, Aldrich and Sumitomo) generated large initial blanks equivalent to 50 to >200 /xM of DOC in a 200-/xL injection. A sample of the type of support distributed by Sumitomo and used in the Sugimura and Suzuki HTCO system gave an initial instrument blank value equivalent to 90 /xM of DOC, which could only be reduced to 27 di 5 /xM by 100 sequential injections of reoxidized water. DON measurements (by persulfate or HTCO) approaching the 40 /xM levels routinely reported by Suzuki et al (1985) and Sugimura and Suzuki (1988) were not later measured (e.g. Walsh, 1989; Hansell, 1993; Koike and Tupas, 1993; Karl et al, 1993). Subsequent to these findings, and to a report by Tanoue (1992) of much lower DOC and DON concentrations obtained in the western North Pacific with an improved HTCO analyzer, Suzuki (1993) retracted the data presented in both his 1985 and 1988 papers. The reasons for the unusually high and closely correlated DON and DOC concentrations appearing in the two retracted papers were not completely clear, although inappropriate attention to instrument blanks was apparently a major problem (Suzuki, 1993). Essentially none of the concentration,

Why Dissolved Organics Matter

23

elemental composition or size distribution results published in the two Suzuki papers has been subsequently confirmed.

VI. WHY DISSOLVED ORGANICS MATTER In view of the huge community response to "New" DOC and the subsequent spate of research on DOM that is described in the following chapters, it is interesting to evaluate why this high level of interest and productivity has been sustained through what could have been a discouraging setback brought about by the Suzuki retractions. In the case of DOM research, much of the reason for this continued level of research activity is traceable to major advances in parallel fields. One of these allied developments has been steadily increasing interest in the global carbon cycle, especially as it relates to greenhouse gases such as CO2 and associated climate change. The birth of this movement can be traced to Svante Arrhenius (1896), who pointed out that humans are increasing atmospheric CO2 concentrations by burning fossil fuels. Arrhenius made the remarkable estimate that a doubling of atmospheric CO2 concentration would lead to a 5-6°C increase in the average temperature of the Earth's surface. Over 50 years later, the reality of increasing atmospheric CO2 concentrations was demonstrated by direct measurements (e.g.. Keeling, 1973; Keeling et al, 1995) and the magnitude of the Arrhenius temperature projection was supported by numeric global climate models (Houghton et al, 1996). Both the great size and potential dynamics of the ocean DOM pool have brought it within the focus of global cycle research (Williams and Druffel, 1988; Toggweiler, 1989b, Hedges, 1992). Because the amounts of carbon in oceanic DOM (--700 x 10^^ g) and atmospheric CO2 (^750 x 10^^ g) are similar (Siegenthaler and Sarmiento, 1993), net oxidation of only 1 % of the seawater DOM pool within 1 year would be sufficient to generate a CO2 flux larger than that produced annually by fossil fuel combustion. Concentration differences of this magnitude would be extremely difficult to identify due to the current limits of analytical precision and the heterogeneous distributions of DOC in the ocean. It is not surprising, therefore, that the Sugimura and Suzuki 1988 report of roughly twice as much total DOC in the ocean as was previously measured gained the immediate attention of the oceanographic conmiunity. In addition to greatly raising the global stakes for budgeting actively cycling organic carbon, this report placed three to six times more DOM in the surface ocean where the bulk of this uncharacterized material appeared to be biologically active on time scales of years to decades (Toggweiler, 1989b). This apparent increase in the organic acid component of seawater was also of substantial interest as a potential explanation for the discrepancy between total CO2 measurements in seawater by potentiometric versus manometric methods

24

John I. Hedges

(Bradshaw and Brewer, 1988). Quantitative constraints on organic matter cycling are particularly difficult in the physically and biologically active upper surface ocean (Quay, 1997), where DOC versus POC exports are difficult to distinguish and DOM photodegradation can accompany photosynthesis. Given the susceptibility of DOC to photolysis (Kieber et ai, 1989; Vodacek et al, 1997) and subsequently biodegradation (Benner and Biddanda, 1998), as well as the rapid increases in UV irradiation of the deeply mixed hub for global thermohaline circulation in the Southern Ocean (Solomon, 1999), an assumption that the contemporary oceanic DOM pool is at steady state seems questionable. Another development that has greatly increased interest in DOM distributions and dynamics over the past three decades has been growing recognition that dissolved organic substrates are important intermediates in rapid cycling of bioactive elements within the ocean (Pomeroy, 1974; Azam and Hodson, 1977). This "microbial loop" from DOM to bacteria, to protists and zooplankton became evident from measurements of heterotrophic bacterial production that typically demanded 20-40% of the local average carbon fixation rate (Azam and Fuhrman, 1984). The only means of supplying such a large a flux of nutrients is by rapid cycling of DOM released by a variety of processes including phytoplankton exudation, viral lysis, and protozoan and zooplankton grazing (Jumars et ai, 1989; Nagata, 2000). Given a global net primary production of ~50 x 10^^ g C/year, the microbial loop would appear to pass the DOM equivalent of at least 10-20 x 10^^ g C/year. If applied uniformly throughout the ocean, this respiration flux could turn over the entire marine DOC pool in less than 100 years, compared to its ^"^C-based "age" of thousands of years. The reason for this discrepancy, of course, is that a small fraction of seawater DOC is recycled biologically in the surface ocean at an exceedingly rapid rate. The "survivor" molecules left behind to accumulate in the DOM pool must nevertheless be subjected to continuous and severe bacterial pressure. Thus, any chink in their molecular armor, such as might be imparted by photolysis (Benner and Biddanda, 1998), abiotic chemical oxidation (Sunda and Kieber, 1994), or physical transformation into gels (Chin et ai, 1998) might lead to rapid and efficient remineralization to CO2. Conversely, the chemical and conformational characteristics of those organic substances that can withstand such concerted attacks for thousands of years in the ocean DOM pool may carry the molecular Rosetta stone for deciphering the degradation mechanisms responsible for recycling the other 99.9% of global primary production (Hedges, 1992). Thus, we have much to learn from both the fast- and the slow-cycling components of ocean DOM. There are many other reasons for continued and growing interest in the forms and reactions of seawater DOM. For example, some molecules dissolved in seawater strongly complex trace metal ions, greatly affecting their bioavailability and toxicity (Buffle, 1988; Kozelka and Bruland, 1998). Dissolved organic molecules also can affect the surface properties of minerals (Stumm, 1992), act as aquatic

Why Dissolved Organics Matter

25

telemediators (Gauthier and Aubert, 1981), and change the spectral properties of seawater (Whitehead and Vemet, 2000). Recent demonstrations that the lignin components of seawater vary in composition and concentration with geographic source (Opsahl and Benner, 1997; Opsahl et ai, 1999) and photodegradation history (Opsahl and Benner, 1998) point toward a future where dissolved organic molecules will provide detailed information about the origins and physical histories of their parent waters. Ultimately, however, the biogeochemical usefulness of any class of chemical tracers is limited by its structural diversity (Blumer, 1976) and the range of its sources, input functions, and chemical reactivities (Middelburg, 1989). It is clear, therefore, that the information content of organic molecules, which also carry imbedded stable isotopic signatures and radiochemical clocks, is unsurpassed by any other seawater component. The 10^^ diverse organic molecules dissolved in every milliliter of seawater are the only constituents whose stored information approaches the richness needed to understand where that water has been and what has happened within it over time. The future of oceanographic research belongs in large part to those who can learn to read these molecular messages.

VII. WHAT DID WE LEARN? Setbacks can be useful learning experiences. The following chapters are testimony that the oceanographic community not only persisted through the "New" DOC experience, but also gained substantially from it in numerous ways over the past decade. First of all, HT(C)0 (now often used without oxidation catalyst) analyzers were fundamentally a great idea and are becoming the instruments of choice for analyses of DOC and DON in laboratories and aboard ships. With appropriate attention to blanks and sample handling, the precision and accuracy of DOC analysis by HT(C)0 have improved to the point that meaningful comparisons among deep-ocean waters and within surface ocean time series have become feasible. In addition, the minimal volume requirements of these analyzers have opened the door to multiple analyses of limited volumes of water from such sources as sediments and experimental incubations. It seems feasible that HT(C)0 analyzers will soon be employed for rapid analysis of the stable carbon isotope composition of DOC in individual samples or used in batch mode to obtain sufficient carbon for ^"^C analysis by accelerator mass spectrometry. Second, we (should) also have come to appreciate the immense importance of carefully developing and rigorously testing new analytical methods. Although fresh concepts will continue to be the primary means of advancement in oceanographic research, it is clear from this retrospective that new perceptions almost always ride the back of improved methods for DOM isolation (e.g., dipped prisms, hydrophobic sorption, and tangential-flow ultrafiltration) and characterization (e.g., stable isotopes, NMR and sensitive molecular analyses). Along with the power to make

26

John I. Hedges

such major advances, however, comes the responsibiUty to adequately test and describe new analytical methods, pointing out their weaknesses as well as their advantages. Editors and reviewers of papers describing new analytical methods also carry the burden of protecting the scientific community (and authors) from published oversights that can propagate for years to great disadvantage. Given the reality that it is often the makers of analytical tools, rather than the wielders, who pace modem scientific advances, this rare skill seems underappreciated overall. Many journals covering the aquatic sciences in fact exclude or strongly discourage "analytical papers," even when the research they describe clearly has been developed specifically to attack biogeochemical research problems. Funding agencies and peer reviewers are often reluctant to fund strictly analytical projects or proposals that involve innovative measurement techniques. Analytical chemistry departments in many universities are presently being dismantled or recombined into topical entities that no longer emphasize or adequately train students in the basics of sound analytical methods. Although the current trend by oceanographers toward broad interests and general skills is healthy, a critical mass of analytically oriented investigators with sufficient chemical understanding to imagine the potential pitfalls involved with measurements of trace organic substances in an ocean of salt will always be required. Finally, the Seattle DOC/DON Workshop and subsequent efforts on the part of Jon Sharp and many other oceanographers demonstrated the tremendous logistical advantage of readily available reference samples and the power of communitywide efforts focused on a shared challenge. The crucial demonstration of a problem in DOC and DON analyses in the early 1990s (Fig. 11) came directly from painstaking analyses by the more than 25 different research groups that participated voluntarily in the Seattle Workshop. Fittingly, the diagnostic offsets in this data set also indicated an experimental path toward resolution that proved to be fruitful. The DOC/DON community also is notable as one of the few oceanographic guilds with an organic orientation that successfully has estabhshed a system for providing widely available reference samples (of low-DOC and deep-ocean water) for blank testing and comparisons among (and within) individual labs. Notably, both actions have involved close collaborations in planning and execution within the marine scientific community and their funding organizations. That these combined efforts have borne such extensive scientific fruit over the past decade (see following chapters) bodes well for future research on the fascinating topic of dissolved organic molecules in the ocean.

ACKNOWLEDGMENTS I thank the editors of this book for their encouragement and guidance. This manuscript benefited greatly from reviews by Ron Benner, Ellen Druffel, John Farrington, Michael Peterson, Jon Sharp, and Kenia Whitehead.

Why Dissolved Organics Matter

27

REFERENCES Anderson, L. A. (1995). On the hydrogen and oxygen content of marine phytoplankton. Deep-Sea Res. 42,1675-1680. Arrhenius, S. (1896). On the influence of carbonic acid in the air upon the temperature of the ground. Phil Mag. 41, 237-276. Azam, R, and Fuhrman, J. A. (1984). Measurement of bacterioplankton growth in the sea and its regulation by environmental conditions. In "Heterotrophic Activity in the Sea" (J. E. Hobbie and P. J. LeB. WiUiams, Eds.), pp. 179-196. Plenum Press, New York. Azam, P., and Hodson, R. E. (1977). Size distribution and activity of marine microheterotrophs. Limnol. Oceanogr. 22,492-501. Bada, J. L., and Lee, C (1977). Decomposition and alteration of organic compounds dissolved in seawater. Mar. Chem. 5,523-534. Baier, R. E. (1972). Organic films on natural bodies of water: Their retrieval, identification and modes of elimination. /. Geophys. Res. 77,5062-5075. Baier, R. E., Goupil, D. W., Perlmutter, S., and King, R. (1974). Dominant chemical composition of sea surface films, natural slicks and foams. /. Rech. Atmos. 8, 571-600. Barber, R. T. (1968). Dissolved organic carbon from deep waters resists microbial oxidation. Nature 220,274-275. Benner, R., and Biddanda, B. (1998). Photochemical transformations of surface and deep marine dissolved organic mater: Effects on bacterial growth. Limnol. Oceanogr. 43,1373-1378. Benner, R., Pakulski, J. D., McCarthy, M., Hedges, J. I., and Hatcher, R G. (1992). Bulk chemical characteristics of dissolved organic matter in the ocean. Science 255,1561-1564. Benner, R., and Strom, M. (1993). A critical evaluation of the analytical blank associated with DOC measurements by high-temperature catalytic oxidation. Mar. Chem. 41,153-160. Blumer, M. (1976). Polycyclic aromatic compounds in nature. Sci. Am. 234, 3 4 ^ 5 . Bradshaw, A. L., and Brewer, P. 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. Broecker, W. S., and Peng, T-H. (1982). "Tracers in the Sea." Columbia University, Palasades, NY. Buffle, J. (1988). "Complexation Reactions in Aquatic Systems: An Analytical Approach." Ellis Horwood, Chichester. Carlson, D. J., Brann, M. L., Mague, T. H., and Mayer, L. M. (1985). Molecular weight distribution of dissolved organic materials in seawater determined by ultrafiltration: A re-examination. Mar Chem. 55,155-171. Chanu, J. (1959). Extraction de la substance jaune dans les eaux cotieres. Rev. Opt. Theor Instrum. 38, 569-572. Chen, R. P., and Bada, J. L. (1989). Seawater and porewater fluorescence in the Santa Barbara Basin. Geophys. Res. Lett. 16, 687-690. Chin, W-C, Orellana, M. V., and Verdugo, P. (1998). Spontaneous assembly of marine dissolved organic matter into polymer gels. Nature 391, 568-572. Craig, H. (1971). The deep metabohsm: Oxygen consumption in abyssal ocean water. J. Geophys. Res. 76,5078-5086. Degens, E. T. (1970). The molecular nature of nitrogenous compounds in sea water and recent marine sediments. In "Organic Matter in Natural Waters" (D. W. Hood, Ed.), pp. 77-106. Institute of Marine Science, Fairbanks, Alaska. Donard, O. F. X., Lamotte, M., Belin, C , and Ewald, M. (1989). High-sensitivity fluorescence spectroscopy of Mediterranean waters using a conventional or pulsed laser excitation source. Mar Chem. 27,111-136.

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Why Dissolved Organics Matter Meyers-Schulte, K. J., and Hedges, J. I. (1986). Molecular evidence for a terrestrial component of organic matter dissolved in ocean water. Nature 321,61-63. Middelburg, J. J. (1989). A simple rate model for organic matter decomposition in marine sediments. Geochim. Cosmochim. Acta 53, 1577-1581. Momzikoff, A., Santus, R., and Giraud, M. (1983). A study of the photosensitizing properties of seawater. Mar. Chem. 12, 1-14. Mopper, K:, Dawson, R., Liebezeit, G., and Hansen, H-R (1980). Borate complex ion exchange chromatography with fluorimetric detection for determination of saccharides. Anal. Chem. 52, 20182022. Mopper, K., and Stahovec, W. L. (1986). Sources and sinks of low molecular weight organic carbonyl compounds in seawater. Mar Chem. 19, 305-321. Morris, R. J. (1974). Lipid composition of surface films and zooplankton from the eastern Mediterranean. Mar Pollut. Bull. 5, 105-109. 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. 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. Ogawa, H., and Ogura, N. (1992). Comparison of two methods for measuring dissolved organic carbon in sea water. Nature 356, 696-698. Ogura, N. (1975). Further studies on decomposition of dissolved organic matter in coastal seawater. Mar Biol. 31, lOl-ni. Ogura, N. (1977). High molecular weight organic matter in seawater. Mar Chem. 5, 535-549. Ogura, N. (1970). The relation between dissolved organic carbon and apparent oxygen utilization in the Western North Pacific. Deep-Sea Res. 17, 221-231. Opsahl, S., and Benner, R. (1997). Distribution and cycling of terrigenous dissolved organic matter in the ocean. Nature 386,480^82. 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. Plunkett, M. A., and Rakestraw, N. W. (1955). Dissolved organic matter in the sea. Deep-Sea Res. 3, 12-14. Pomeroy, L. R. (1974). The ocean's food web, a changing paradigm. Bioscience 24,499-504. Putter, A. (1909). Die Emahrung der Wassertiere und der Stoffhaushalt der Gewasser. Jena, 1909. Quay, P. (1997). Was a carbon balance measured in the equatorial Pacific during JGOFS? Deep-Sea Res. 44,1765-1781. Raben, E. (1910). 1st organischen Kohlenstoff in nennenswerter Menge im Meerwasser gelost vorhanden? Wiss. Meeresunt. Abt. Kiel NF11, 111-117. Redfield, A. C , Ketchum, B. H., and Richards, F. A. (1963). The influence of organisms on the composition of sea water. In "The Sea" (M. N. Hill, Ed.), pp. 26-77. Interscience, New York. Riley, J. P., and Taylor, D. (1969). The analytical concentration of traces of dissolved organic materials from sea water with Amberlite XAD-1 resin. Anal. Chim. Acta 46, 307-309. Sakugawa, H., and Handa, N. (1985). Isolation and chemical characterization of dissolved and particulate polysaccharides in Mikawa Bay. Geochim. Cosmochim. Acta 49, 1185-1193. Sakugawa, H., Handa, N., and Ohta, K. (1985). Isolation and characterization of low molecular weight carbohydrates dissolved in seawater. Mar Chem. 17, 341-362. Sharp, J. H. (1973a). Size classes of organic carbon in seawater. Limnol. Oceanogr 18,441-447. Sharp, J. H. (1973b). Total organic carbon in seawater—Comparison of measurements using persulfate oxidation, and high temperature combustion. Mar Chem. 1, 211-229.

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Vodacek, A., Blough, N. V., DeGrandpre, M. D., Peltzer, E. T., and Nelson, R. K. (1997). Seasonal variation of CDOM and DOC in the Middle Atlantic Bight: Terrestrial inputs and photooxidation. Limnol. Oceanogr. 42, 674-686. Waksman, S. A., and Carey, C. L. (1935). Decomposition of organic matter in sea water by bacteria. /. Bacteriol 29,531-543. Walsh, T. W. (1989). Total dissolved nitrogen in seawater: A new high-temperature combustion method and a comparison with photo-oxidation. Mar. Chem. 26, 295-311. Whitehead, K., and Vemet, M. (2000). Influence of mycosporine-like amino acids (MAAs) on the UV absorption of particulate and dissolved organic matter in La JoUa Bay. Limnol Oceanogr. 45, 1788-1796. Willey, J. D., and Atkinson, L. P. (1982). Natural fluorescence as a tracer for distinguishing between piedmont and coastal plain river water in the nearshore waters of Georgia and North Carolina. Estuarine Coastal Shelf ScL 14,49-59. Wilhams, P. J. le B. (1975). Biological and chemical aspects of dissolved organic matter in sea water. In "Chemical Oceanography" (J. P. Riley and G. Skirrow, Eds.), Vol. 2, pp. 301-363. Academic Press, New York. Wilhams, P. M. (1968a). Organic and inorganic constituents of the Amazon River. Nature 218,937-938. Wilhams, P. M. (1968b). Stable carbon isotopes in the dissolved organic matter of the sea. Nature 219, 152-153. Williams, P. M. (1969a). The wet oxidation of organic matter in seawater. Limnol. Oceanogr 14, 292-297. Williams, P. M. (1969b). The determination of dissolved organic carbon in seawater: A comparison of two methods. Limnol Oceanogr 14, 297-298. Wilhams, P. M. (1971). The distribution and cycling of organic matter in the ocean. In "Organic Compounds in the Aquatic Environment" (S. D. Faust and J. V. Hunter, Eds.), pp. 145-163. Dekker, New York. Williams, P. M., and Carlucci, A. F. (1976). Bacterial utihsation of organic matter in the deep sea. A^fl^Mr^ 262, 810-811. Wilhams, P. M., and Druffel, E. R. M. (1987). Radiocarbon is dissolved organic matter in the central North Pacific Ocean. Nature 339,246-248. Wilhams, P. M., and Druffel, E. R. M. (1988). Dissolved organic matter in the ocean: comments on a controversy. Oceanography 1, 14-17. Williams, P. M., and Gordon, L. I. (1970). Carbon-13:carbon-12 ratios in dissolved and particulate organic matter in the sea. Deep-Sea Res. 17,19-27. Williams, P. M., Oeschger, H., and Kinney, P. (1969). Natural radiocarbon activity of dissolved organic carbon in the Northeast Pacific Ocean. Nature 224, 256-258. Wilson, D. P., and Armstrong, F. A. (1952). Further experiments on biological differences between natural sea waters. /. Mar Biol Assn. UK 31, 335-349. Wilson, M. A., Gillam, A. H., and Collins, P. J. (1983). Analysis of the structure of dissolved marine humic substances and their phytoplanktonic precursors by ^H and ^^C nuclear magnetic resonance. Chem. Geol 40, 187-201. Wilson, M. A., Philp, R. P, Gillam, A. H., Gilbert, T. D., and Tate, K. R. (1983). Comparison of the structures of humic substances from aquatic and terrestrial sources by pyrolysis gas chromatographymass spectrometry. Geochim. Cosmochim. Acta 47,497-502. Zepp, R. G., Schlotzhauer, P. F., and Sink, R. M. (1985). Photosensitized transformations involving electronic energy transfer in natural waters: Role of humic substances. Environ. ScL Technol 19, 74-81.

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Chapter 2

Analytical Methods for Total DOM Pools Jonathan H. Sharp Graduate College of Marine Studies, University of Delaware, Lewes, Delaware I. Introduction II. Dissolved Organic Carbon Analysis A. Historical Perspective B. The Analytical Problem C. Small Group Methods Comparisons D. Broad Community Methods Comparisons III. Dissolved Organic Nitrogen Analysis A. Historical Perspective and the Analytical Problem B. Small Group Methods Comparisons

C. Current and Future Broad Community Methods Comparison IV. Dissolved Organic Phosphorus Analysis A. Historical Perspective and the Analytical Problem B. Small Group Methods Comparisons V. Multielemental Methods VI. The Limits of Elemental Analyses VII. The Need for Continual use of Reference Materials References

I. INTRODUCTION For over a century, there has been interest in quantifying the pool of dissolved organic matter (DOM) in the sea and in other aquatic environments. Some of the early efforts included attempts to quantify "humic substances" and to measure colored material, "gelbstoffe" (Duursma, 1965). Specific organic compounds have been measured and attempts have been made to understand origins and fates of Biogeochemistry of Marine Dissolved Organic Matter Copyright 2002, Elsevier Science (USA). All rights reserved.

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these compounds for decades (Duursma, 1965; Williams, 1975; Menzel, 1974). Analysis of specific compounds and dynamics of organic matter in the sea are discussed in several other chapters of this book. While that type of research is possibly more interesting and fruitful, depending upon the questions asked, more energy has been expended overall to quantify the gross amount of DOM. For much of the past century, the usual currency for quantifying the DOM has been dissolved organic carbon (DOC). A general rule of thumb, not much used today, was that the weight of the total organic matter pool is two times that of the measured organic carbon (Krogh, 1934). Thus, one of the main reasons for measuring DOC has been to quantify the total DOM. Fewer measurements have also been made over the years of dissolved organic nitrogen (DON) and dissolved organic phosphorus (DOP). It is unfortunate that sometimes DOC, DON, and DOP are viewed as specific and separate entities. It is probably useful to recognize that these are three subsets of the same DOM. Measurement of them is quantitative, but not really qualitative in terms of characterizing the DOM pool. They take on separate importance when considering the total pools of carbon, nitrogen, and phosphorus in the sea. Measurements of DOC, DON, and DOP have often been performed routinely as minor side parameters in biological or multidisciplinary studies; until recently, these analyses were downplayed or ignored by much of the geochemical community. In this chapter, the focus is on "dissolved" organic matter with the assumption that particles can be removed by filtration through a microporous filter. Such filtration, essential in nearshore waters, is often omitted in oceanic waters where the total organic matter is almost all "dissolved"; e.g., TOC ^ DOC. With the high DOC and DON concentrations reported by Suzuki (Suzuki et al, 1985; Sugimura and Suzuki, 1988), measurements of DOM took on new importance in relation to the global carbon cycle (Toggweiler, 1989). These high concentrations and odd depth distributions also required rethinking of nitrogen and phosphorus cycles as well as carbon cycles (Williams and Druffel, 1988; Jackson, 1988; Toggweiler, 1992). The "controversy" on DOM measurements led to the Seattle workshop (Williams, 1992) that has guided much effort since then (Sharp, 1997). The increased international interest on measurement of DOM makes accurate analyses critical. Prior to this new interest, there were many small "internal" methods checks made to verify accuracy of methods used routinely. As will be discussed below, such efforts have not been very successful; larger, broad-community, methods comparisons ultimately can prove to be more effective in developing community analytical accuracy. Also, discussed below is the opinion of this author that strict continuing use of reference materials is essential for community accuracy. In this chapter, there will be decreasing emphasis from DOC to DON to DOP, reflective of the efforts both in performing routine measurements and in establishing accurate methods. Although routine measurements have been made for decades, the groundswell of interest has propelled DOC analysis into the limelight for the

Analytical Methods for Total DOM Pools past 15 years. Currently, there is more interest in the measurement of DON than in the recent past and an international methods comparison is ongoing through my laboratory. DOP has been more of a poor relative compared to DOC, but some small internal comparisons have been made recently, which will be discussed. Analyses of DOC, DON, and DOP are considered separately below with some historical perspective and discussion of both the small "internal" and the "broadcommunity" methods comparisons. Analytical problems in analysis of DOC are distinctly different from those of DON and DOP; so there is a brief discussion on the analytical problem in each section below.

11. DISSOLVED ORGANIC CARBON ANALYSIS A. HISTORICAL PERSPECTIVE From the earliest published measurements of DOM in seawater (Natterer, 1892), disagreement arose on the accuracy of the quantification and biological significance of the DOM pool. With methods that were soon criticized. Putter reported high concentrations of DOC and suggested that the high concentrations which originated from plankton blooms served as a major food for invertebrates (Putter, 1909). His work was questioned for accuracy of measurement of DOC and DON; ultimately, better methodology discounted his work (Krogh and Keys, 1934). The methods used for the earlier research and all other methods through the 1950s employed some form of wet chemical oxidation (WCO) and colorimetric or gravimetric determination of the resultant oxidation product, CO2. In the late 1950s, Soviet scientists started using high-temperature combustion (HTC) for oxidation of dried samples followed by colorimetric analysis of CO2 (see Skopintsev et al, 1966). The first electronic measure of the CO2 came with the persulfate oxidation (PO) method that culminated in nondispersive infrared analysis (Wilson, 1961). This method was soon popularized as a standard method (Menzel and Vaccaro, 1964). A similar PO method that did not catch on used gas chromatography for the final CO2 determination (Fredericks and Hood, 1965). The PO method with infrared detection became the one most used by the late 1960s. Also in the 1960s, UV oxidation was established as another method for DOC analysis (Armstrong et al, 1966; Williams, 1969); but was never established as a broadly used DOC method.

B. THE ANALYTICAL PROBLEM The quantitative estimate of DOC concentration has almost always been done by conversion of all organic carbon to CO2. Prior to making the conversion.

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Jonathan H. Sharp Acidify sampie and sparge to remove dissolved inorganic carbon.

UV oxidation to convert all DOC to CO2.

Persulfate oxidation to convert all DOC to CO2.

Colorimetric measurement of CO2. All resultant CO2 represents original DOC. Figure 1

High temperature oxidation to convert all DOC to CO2.

Measurement of CO2 in nondispersive infrafed analyzer. All resultant CO2 represents original DOC.

Schematic diagram of procedural steps in commonly used DOC methods.

the dissolved inorganic carbon (DIC) must be removed from the sample, the acid-sparging step (Fig. 1). The procedure used is to acidify the sample, lowering the pH to below 2. This shifts the carbonate equilibrium so that almost all of the DIC is in the form of dissolved CO2. Bubbling this acidified sample with a gas free of CO2 will quickly remove the entire DIC pool. Actually, the removal of CO2 is asymptotic and a tiny amount remains. Experimentation has shown that the residual CO2 can easily be kept below 0.01% of the original (e.g., Sharp, 1977). After removal of the DIC, rapid and complete conversion of the DOC to CO2 requires strong oxidation (Fig. 1). Early DOC methods used exposure to strong acids or other oxidants. There is the inherent error that any easily degraded organic molecules or those that could be made volatile under mildly acidic conditions will be lost in the acid-sparging step. It has always been assumed that the acid-volatile losses are small and it has been demonstrated that the losses are on the order 100 Dalton, indicating that combined forms of DOM that pass an ultrafiltration membrane with a 1000-Dalton molecular weight cutoff consist of 1000 m) Ocean Chemical characteristics

Surface ocean

Deep ocean

Reference

Bulk composition DOC (/xM) DON (ixM) DOP iixM) DOC:DON DOC:DOP Carbohydrates (/xM C glucose equivalent)

60-90 3.5-7.5 0.1-0.4 9-18 180-570 10-25

35^5 1.5-3.0 0.02-0.15 9-18 300-600 5-10

1-6,9-10 6-11 6-12 1-12 1-12 13-14

200-800 200-500 42-94 0.2-0.7

20-170 80-160

15-16 17-18 19-20

Molecular composition Total hydrolyzable neutral sugars (nM) Total hydrolyzable amino acids (nM) Total hydrolyzable amino sugars (nM) Lipids (solvent extractable; nM) Total hydrolyzable neutral sugars (% DOC) Total hydrolyzable amino acids (% DOC) Total hydrolyzable amino sugars (% DOC) Solvent extractable lipids (% DOC) Total (% DOC) Total hydrolyzable amino acids (% DON) Total hydrolyzable amino sugars (% DON) Total (% DON)

2-6 1-3

0.4-0.6 0.3-0.9 3.7-10.5 6-12 0.8-1.7 6.8-13.7

4-9 nd

0.5-2.0 0.8-1.8 0.04-0.07

21

nd

1.3-3.9

4-9

0.2-0.4 4.2-9.4

Note. Some measurements are for unfiltered water samples and thus include small ( 1000 m) marine DOC and DON identified as specific biomolecules (total hydrolyzable neutral sugars, total hydrolyzable amino sugars, and total hydrolyzable amino acids). The uncharacterized fraction is DOC and DON not accounted for in these specific biomolecules.

at the molecular level. For comparison, the percentage of total carbon comprised by these biochemicals is ^6.6% in surface water DOM and --2.0% in deep water DOM (Table 1). This comparison suggests that DOM in the deep ocean is the most diagenetically altered, as well as abundant, form of organic matter in the sea.

C. CHEMICAL COMPOSITION OF ISOLATED FRACTIONS OF

DOM

The concentration and isolation of DOM from seawater is advantageous for its chemical and isotopic characterization. As discussed in the Introduction, a variety

72

Ronald Benner

Phytoplankton

Sediment trap

U • •

Surface sediment

Surface DOM

Amino acids Lipids Neutral sugars

a

Deep DOM

20

40

60

80

100

Percent characterized organic carbon

Figure 3 Contributions of total hydrolyzable amino acids, total hydrolyzable neutral sugars, and lipids to the organic carbon content of net phytoplankton, sinking particles collected in deep water sediments traps, surface sediments from the deep sea, and DOM from the surface ( 1000 m) ocean. Plankton, sediment trap, and surface sediment data are from Wakeham et al. (1997).

of solid-phase extraction techniques have been used to isolate the operationally defined humic fraction of DOM. Data on the chemical composition of humic substances isolated with the most commonly used sorbents, XAD-2 and XAD-8 resins, are presented in Table II. The bulk C:N ratios for dissolved humic substances isolated from surface and deep water range from 36 to 57, indicating that humic substances comprise a highly N-depleted fraction of DOM. Organic nitrogen is often charged and relatively polar and is therefore depleted in these relatively hydrophobic fractions of DOM. Nuclear magnetic resonance (NMR) spectroscopy is a very useful approach for determining the bulk chemical structure of natural organic matter, and its application to isolated fractions of DOM has provided novel insights about the structural relationships of major bioelements in DOM, including C, H, N, and P. Solid-state ^^C NMR analyses of humic substances from the Pacific Ocean indicate that there is no recognizable compositional difference between surface and deep-water counterparts (Table II). This similarity in structure suggests that humic substances comprise a relatively refractory fraction of DOM. On average.

73

Chemical Composition and Reactivity Table II Average Chemical Characteristics of Dissolved Organic Matter (DOM) Isolated from Surface and Deep Ocean Waters Chemical characteristics

Surface ocean

Deep ocean

Reference

Humic substances (% DOC) Humic substances (C/N atom) Humic substances (% C-C) Humic substances (% C-0, 0-C-O) Humic substances (% C ^ C ) Humic substances (% COO,CNO) Humic substances (% C = 0 ) Total hydrolyzable amino acids (%HS-DOC) Total hydrolyzable amino acids (%HS-DON) HMW DOM (% DOC) HMW DOM (C/N atom) HMW DOM (% C-C) HMW DOM (% C-O, 0-C-O) HMW DOM (% C = C ) HMW DOM (% COO, CNO) HMW DOM (% C = 0 ) Total hydrolyzable neutral sugars (%HMW DOC) Total hydrolyzable amino acids (%HMW DOC) Total hydrolyzable amino sugars (%HMW DOC) Total % DOC Total hydrolyzable amino acids (%HMW DON) Total hydrolyzable amino sugars (%HMW DON) Total % DON

5-25 36-56

15-25 39-57

1-5 3-4 4 4 4 4 4

5-11 25-40 15-18

20-25 18-20

6-13

1.5-2.5

1.3-2.6 10.3-19.6 17-29 3.6-7.1 20.6-36.1

0.5-0.6 5.0-7.1 12-28 1.4-1.9 13.4-29.9

44 19 19 15 3 nd

25 54 5 13 3

3-4

46 17 19 15 3 nd 5-9 30 28 21 16 5

3>-4

6-7

8-10 8,10 8,11 8,11 8,11 8,11 8,11 12-13

12 14

12-13

12 14 12

Note. Humic substances are operationally defined as DOM that is retained on XAD-2 or XAD-8 resins under acidic conditions. High-molecular-weight (HMW) DOM is operationally defined as DOM that is retained by an ultrafiltration membrane with a ~ l - n m pore size and 1000-Dalton molecular weight cutoff. Cross-polarized magic angle spinrung (CPMAS) ^^C nuclear magnetic resonance spectroscopy was used for carbon functional group analyses. References. 1 Gagosian and Stuermer 1977; 2, Harvey et al. (1983); 3, Druffel et al. (1992); 4, Hedges et al. (1992); 5, Ishiwatari (1992); 6, Hubberton et al. (1994); 7, Hubberton et al. (1995); 8, Benner et al (1992); 9, Guo et al. (1995); 10, Benner et al. (1997); 11, McCarthy et al. (1993); 12, McCarthy et al. (1996); 13, Skoog and Benner (1997); 14, Benner and Kaiser unpublished.

humic substances are dominated (^45% of C) by hydrogen-substituted alkyl C, such as methylene groups (Table II; Fig. 4). Oxygen-substituted alkyl C, such as carbohydrate-C, accounts for slightly less than 20% of the C in humic substances. Unsaturated C, such as aromatic-C, accounts for '^20% of the C. About 15% of the C resides in carboxyl and amide groups, although the low

74

Ronald Benner c=o

c=o

coo, CNO

c-c

COO. CNO

c-c

c=c

c=c

C-0, 0-C-O c-0, 0-C-O HMW DOM

Humic substances

Figure 4 The distribution of major functional groups of carbon in high-molecular-weight (HMW) DOM and dissolved humic substances from the surface ocean. The functional group assignments are based on analyses by solid-state cross-polarized magic angle spinning (CPMAS) ^^C NMR.

N content of humic substances indicates that most of this is carboxyl-C (salt and ester). A low percentage (3%) of the C resides in carbonyl compounds (ketones and aldehydes). There have been relatively few molecular-level analyses of conmion biochemicals in dissolved humic substances isolated from seawater. Analyses of amino acids (THAA) indicates that they account for 5-11% of the N in dissolved humic substances (Table II), a value that is similar to the percentage of amino-acid N in DOM (Table I). It is important to recognize that the same common biochemicals found in DOM are often found in dissolved humic substances. Thus, molecularlevel analyses of seawater DOM include biochemicals that reside in the dissolved "humic substances" fraction. As discussed earlier, solid-phase extraction isolates DOM principally on the basis of chemical properties, whereas ultrafiltration isolates DOM principally on the basis of physical properties. Given this fundamental difference in isolation mechanisms, it is particularly interesting to compare the chemical compositions of these fractions of DOM. The C:N ratio of HMW DOM isolated using ultrafiltration ranges from 15 to 20 and is highly enriched in N relative to its humic substances counterpart (Table II). Another striking difference between humic substances and ultrafiltered HMW DOM is in the compositional variability between DOM isolated from surface and deep water. Humic substances are compositionally invariant between the surface and deep ocean, indicating that they are less reactive components of DOM. In contrast, HMW DOM from surface water is highly enriched in oxygen-substituted alkyl C (C-0, 0-C-O) relative to HMW DOM from deep water (Table II). These functional groups are commonly found in carbohydrates, and the C - 0 : 0 - C - 0 ratio falls in the range of 4-5:1 (Benner et ai, 1992), which is

Chemical Composition and Reactivity

75

characteristic of common carbohydrates (i.e., pentoses and hexoses). Slightly over 50% of the C in HMW DOM from surface water is in carbohydrates, whereas only half that percentage is found in HMW DOM from deep water (Table II). Half of all C in surface water HMW DOM is found in carbohydrate-like structures (Fig. 4), indicating that oligo- and polysaccharides are abundant. The large reduction in the percentage of carbohydrate-C in deep-water HMW DOM suggests that oligo- and polysaccharides are reactive components that are largely produced and consumed in the upper ocean. About 25-30% of C in surface and deep DOM resides in substituted alkyl C, which is considerably less than is found in dissolved humic substances (Table II; Fig. 4). Only ~ 5 % of the C in surface DOM resides in unsaturated-C (C=C) structures, whereas ^20% of the C in deep DOM resides in unsaturated-C structures. The major reduction in the fraction of carbohydrate-C between surface and deep HMW DOM is largely compensated by an increase in the fraction of unsaturated-C in deep HMW DOM. Carboxyl-C and amide-C accounted for 13-16% of HMW DOM. The maximum amide-C contribution can be estimated based on the N content if we assume all N is in amide form (McCarthy et al, 1997). Based on an average C:N ratio of 17, a maximum of ^6% of the C in HMW DOM is in amide form, and a minimum of 6-10% of the C is found in carboxyl (salt and ester) structures. A small fraction (3-5%) of the C in HMW DOM resides in aldehyde and ketone structures. The overlap between the C isolated as humic substances and DOM has not been determined, but the present comparison between these fractions indicates substantial compositional differences. Based of the major difference in the carbohydrate-C in surface water humic substances and HMW DOM as indicated by NMR, it appears that the potential overlap is at most --50%. Overlap between these fractions of DOM could be larger in deep water because there was less compositional distinction between these isolated fractions. Molecular-level analyses indicate that neutral sugars (THNS) account for 613% of the C in surface HMW DOM and 1.5-2.5% of the C in deep HMW DOM (Table II). Amino sugars (THAS) account for an additional 1.3-2.6% of the C in surface HMW DOM and 0.5-0.6% of the C in deep HMW DOM (Table II). The NMR analysis of surface and deep HMW DOM samples indicated that carbohydrates accounted for ^50 and '^25% of the C, respectively. Together, neutral and amino sugars account for 15-20% of the carbohydrates in surface HMW DOM and 5-10% of the carbohydrates in deep HMW DOM indicated by NMR. A much larger fraction (50%) of the total carbohydrates in freshly produced HMW DOM from phytoplankton is characterized as neutral sugars (Biersmith and Benner, 1998). The fraction of total carbohydrates identified as neutral sugars appears to decrease with increasing diagenetic alteration of DOM. Carbohydrates are a diverse class of compounds, and only a subset is currently analyzed at the molecular level. Thus, it is possible that the predominant

76

Ronald Benner

carbohydrates in seawater are those that we do not analyze at the molecular level. Variability in susceptibility to hydrolysis is also a problem in carbohydrate analyses at the molecular level. Optimization of hydrolysis is a balance between conditions that are strong enough to release monomers from hydrolysis-resistant matrices and conditions that do not destroy the molecular identity of the monomers. It is also possible that some of the oxygen-substituted carbon identified by NMR does not reside in carbohydrates. We are unsure of the explanation for this important observation, but it demonstrates the usefulness of combining bulk and molecular analyses in attempting to unravel the mysteries concerning the origins and transformations of DOM. Amino acids (THAA) account for a relatively constant 3 ^ % of the C in surface and deep HMW DOM (Table II). A larger fraction (12-29%) of the N in surface and deep HMW DOM is accounted for as amino acids. Amino sugars account for 3.6-7.1% of the N in surface HMW DOM and 1.4-1.9% of the N in deep HMW DOM. Thus, a total of ^ 1 5 % of the C and 28% of the N in surface HMW DOM resides in neutral sugars, amino sugars, and amino acids, whereas ^ 6 % of the C and ~ 2 1 % of the N in deep HMW DOM resides in these common biochemicals. The percentages of total carbon accounted for as neutral sugars, amino sugars, and amino acids in DOM and HMW DOM from surface and deep water are compared in Fig. 5. Two trends in these data are particularly important. The percentages of carbon accounted for in these biochemicals are greater in HMW DOM relative to DOM, and the respective percentages of carbon are greater in surface water relative to deep water. As discussed earlier, the yields of these conmion biochemicals are useful indicators of the diagenetic history of the organic matter. The higher biochemical yields in surface relative to deep DOM and HMW DOM indicate that surface waters contain organic matter of more recent origin that is less altered and potentially more bioreactive than organic matter in deep water. This interpretation of the biochemical data is consistent with the contrasting average apparent ages of surface and deep DOM (Williams and Druffel, 1987) as well as their relative bioreactivity (Barber, 1968). The higher yields of neutral sugars, amino sugars, and amino acids in HMW DOM relative to DOM indicate that the HMW components of DOM are less diagenetically altered and potentially more bioreactive than the LMW components. Direct comparisons of the bioreactivity of the HMW and LMW components of DOM in a variety of marine environments also indicates that the bulk of HMW DOM is more bioreactive than the bulk of LMW DOM (Amon and Benner, 1994, 1996a). These varying lines of evidence lead to the observation that LMW DOM is the most diagenetically altered and least bioreactive fraction of organic matter in seawater. The three groups of biochemicals analyzed in DOM and HMW DOM also appear to have varying susceptibilities to diagenetic alteration. Neutral sugars and amino sugars undergo greater relative losses between surface and deep DOM

u* L 1

77

Chemical Composition and Reactivity

1 — 1 — 1 — 1 — 1 — 1 — 1 —

Surface DOM

Surface HMW DOM

Deep DOM

Deep HMW DOM

L

H B D

10

Amino acids Amino sugars Neutral sugars

15

20

Percent characterized organic carbon

Figure 5 Contributions of total hydrolyzable amino acids, total hydrolyzable neutral sugars, and total hydrolyzable amino sugars to the organic carbon content of DOM and high-molecular-weight (HMW) DOM from the surface (< 100 m) and deep (> 1000 m) ocean.

and HMW DOM than do amino acids, suggesting that most of these combined sugars are more bioreactive than combined amino acids. Seawater HMW DOM has been analyzed by solid state ^^N NMR and ^^P NMR aswellasi^CNMR(Benner^?a/., 1992; McCarthy ^f a/. 1991; Clark etai, 1998), providing a unique perspective on the chemical composition of this fraction of surface and deep HMW DOM (Fig. 6). It is quickly apparent that there is little variability between the ^^N and ^^P spectra of surface and deep HMW DOM, compared to substantial differences between ^^C spectra of surface and deep HMW DOM (as previously discussed). The concentrations of carbon in major functional groups of HMW DOM from surface and deep water are compared in Table III. Assuming the HMW DOM in surface and deep water had similar concentrations and compositions at the time of formation, the difference between the surface and deep HMW DOM compositions is representative of the carbon removed during decomposition. This calculation indicates that 70% of the carbon removed was oxygen-substituted, as occurs in carbohydrates. The C - 0 : 0 - C - 0 ratio of the removed fraction was in the 4-5:1 range expected for common carbohydrates. By this analysis, combined carbohydrates are the most reactive components of HMW DOM.

78

Ronald Benner

l^C-NMR

300

200

100

-100

l^N-NMR

450 " ' " ' * 350 " " ' " 250 '

40

20

«I''''

I'''

0

150

50

r I r y f i » I I I I I ' I '''

-20

-40

Chemical shift (ppm) Figure 6 Solid-state cross-polarized magic angle spinning (CPMAS) ^^C NMR, ^^N NMR, and ^^P NMR spectra of surface and deep high-molecular-weight (HMW) DOM from the surface (< 100 m) and deep (> 1000 m) Pacific ocean. Spectra were taken from Benner et al. (1992), McCarthy et al. (1997), and Kolowith et al. (2001). The asterisks on the ^^N NMR spectra mark the locations of spinning sidebands.

Another interesting feature of the HMW DOM comparison in Table III is the near constant relationship between the C-C and (COO + CNO) functional groups. The C-C:(COO -h CNO) ratio is 1.92,1.87, and 1.97 in surface, deep, and "difference" HMW DOM groups, respectively. Combined amino acids contain all of these functional groups, but comprise a small fraction (3-4%) of the chromatographically

79

Chemical Composition and Reactivity Table III

The Concentrations (/>tM) of Organic carbon and Nitrogen in Surface and Deep High-Molecular-Weight (HMW) DOM and the Concentrations of Organic Carbon Identified in Major Functional Groups by (CPMAS) ^^C Nuclear Magnetic Resonance Spectroscopy HMW DOM

DOC

Surface Deep Difference

21.0 8.1 12.9

DON

ifJLM)

1.25 0.45 0.80

c-c

o-c-o

c=c

COO, CNO

C=0

5.25 2.43 2.82

11.3 2.27 9.03

1.05 1.70 -0.6S

2.73 1.30 1.43

0.63 0.41 0.22

C-0,

Note. The concentrations of DOC and DON are typical values for Pacific Ocean HMW DOM (Benner et al, 1997). The fraction of carbon in major fiinctional groups was taken from Table II.

resolved carbon in surface and deep HMW DOM (Table II). The combined C-C and (COO + CNO) functional groups account for -^40% of the carbon in HMW DOM (Table III). Thus, combined amino acids account for a small fraction of these functional groups in HMW DOM. The molecular "home" of most of these functional groups remains unknown. The^relatively constant ~2:1 ratio of these functional groups in HMW DOM undoubtedly provides clues about its biochemical origin. The constant ratio of these functional groups indicates this material is of average reactivity. The main feature of the ^^N NMR spectra is a resonance centered near a chemical shift of 260 ppm (Fig. 6). This resonance corresponds to the chemical shift for amide nitrogen, which is found in combined amino acids and A/^-acetyl amino sugars (McCarthy et al, 1997). A shoulder is observed on the downfield side of the main resonance at 260 ppm, and this could be indicative of the presence of nitrogen in heterocyclic structures. It is estimated that 75-85% of the nitrogen in HMW DOM is in amide structures. The abundance of amide nitrogen in HMW DOM is perplexing when compared to the amount of measured amino acid (THAA) nitrogen in these samples. Only 8-12% of the amide nitrogen indicated by NMR analyses is identified at the molecular level as hydrolyzable amino acid nitrogen (McCarthy et al, 1997). Proteins and peptides are the most common biopolymers with amide nitrogen. It is possible that a large fraction of the peptide and protein nitrogen in HMW DOM is resistant to hydrolysis, but it is unlikely that hydrolysis-resistant peptides and proteins would be several-fold more abundant than the hydrolyzable fraction that predominates in marine organisms. Biopolymers containing A^-acetyl amino sugars, such as chitin and peptidoglycan, are also abundant in marine plankton. These biopolymers contain amide nitrogen, but analyses of hydrolyzable amino sugars (THAS) in HMW DOM indicate that 80%) of the amide nitrogen in HMW DOM remains unidentified at

80

Ronald Benner

the molecular level. Most (>80%) of the oxygen-substituted carbon (i.e., carbohydrate) indicated by *^C NMR also remains unidentified at the molecular level. When we resolve these differences we will understand fundamental aspects about the sources and transformations of DOM that remain elusive at the present time. Decomposition processes lead to an increase in the fraction of molecularlyuncharacterized organic matter (Hedges et al, 2000). Thus, the above observations suggest a shared diagenetic history for the major biochemicals in DOM. The ^^P NMR spectra of HMW DOM have resonances centered around chemical shifts of 0 and 25 ppm (Fig. 6). The main resonance at 0 ppm is characteristic of phosphate esters, including monoester and diester phosphates (Clark et al, 1998; Kolowith et al, 2001). About 75% of the phosphorus in surface and deep HMW DOM resides in ester structures. Phosphate esters are found in conmion biopolymers, such as membrane phospholipids, RNA, and DNA. The resonance at 25 ppm is characteristic of phosphonate structures, which contain a carbonphosphorus bond. About 25% of the phosphorus in surface and deep DOM resides in phosphonate structures. Phosphonates are much less abundant in biopolymers than phosphate esters, but they are synthesized in phosphonolipids and phosphonoproteins by a variety of marine organisms (Kolowith et al, 2001). The high relative abundance of phosphonates in HMW DOM is somewhat surprising given their relatively low abundance and apparently limited distribution in marine organisms. Analyses by ^H NMR are supportive of ^^C NMR analyses, indicating that surface water HMW DOM is rich in carbohydrates and depleted in aromatic components (Vemonclark etal, 1995; Aluwihare etal, 1997). These studies also noted the presence of acetyl groups in DOM, and it was speculated that the carbohydrates in DOM were predominantly 0-acetyl oligosaccharides (Aluwihare et al, 1997). However, subsequent analyses of these DOM samples by direct temperatureresolved ammonia chemical ionization mass spectrometry did not detect evidence for 0-acetyl oligosaccharides (Boon et al, 1998). The latter authors found evidence for the presence of A/^-acetyl groups and suggested that the acetyl groups indicated by ^ H NMR reside in A/^-acetyl amino sugars rather than O -acetyl oHgosaccharides. The presence of amino sugars in HMW DOM is confirmed by molecularlevel analyses (Table II). As indicated above, A^-acetyl amino sugar polymers in HMW DOM contribute to the amide signal observed in the ^^N NMR spectra.

III. MAJOR TOPICS OF ONGOING AND FUTURE RESEARCH ABOUT THE CYCLING OF DOM A. MECHANISMS OF FORMATION AND REMOVAL OF BlOREFRACTORY D O M Two important observations about the cycling of DOM in the ocean are readily apparent: (1) most of the DOM produced is rapidly consumed, (2) most of the

Chemical Composition and Reactivity DOM remaining in the ocean is resistant to biological utilization. The first observation is apparent from measurements of the growth and respiration rates of microorganisms in surface waters of the ocean. Heterotrophic bacterial production is a large fraction (0.15-0.20) of primary production in the ocean (Ducklow, 2000). Heterotrophic bacteria are the predominant consumers of DOM, and their growth efficiencies are typically less than 30% (del Giorgio and Cole, 2000). Thus, bacteria utilize over 50% of primary production for growth and respiration in the upper ocean. Production and consumption of DOM in the surface ocean are typically tightly coupled, indicating that most DOM produced in the surface ocean is rapidly consumed. A small fraction of the organic matter produced in the surface ocean escapes remineralization and contributes to the slowly cycling, "biorefractory" DOM reservoir. There are numerous observations indicating that most of the DOM in the ocean is resistant to biological utilization. Most DOM (>70%) resides in the deep ocean, and experimental studies of the bioreactivity of deep water DOM (e.g.. Barber, 1968) as well as oceanic surveys of deep water DOC concentrations (Hansell and Carlson, 1998) indicate that this material is of limited bioavailability and reactivity. Average radiocarbon ages of DOM in the deep ocean are 4000-6000 years (WiUiams and Druffel, 1987; Bauer et al, 1992), a further indication of the slow cycling of this material. So, while most of the DOM produced in the ocean is consumed on time scales of hours to weeks, most of the DOM reservoir in the ocean is resistant to biological utilization. There are two key questions in this carbon cycle conundrum. How is biorefractory DOM formed, and how is it removed? Answers to these questions are central to understanding the global carbon cycle as well as carbon cycling in the ocean. Information about the size distribution and chemical composition of DOM presented earlier in this chapter provides some clues about the formation of biorefractory DOM in the ocean. The fraction of LMW DOM, chemically uncharacterized DOM, and biorefractory DOM all increase with depth in the ocean and appear to be related. Enzymatic and photochemical decomposition processes generally lead to decreases in the size and molecular weight of DOM (Amosti, 1998; Smith et al, 1992; Opsahl and Benner, 1998). Direct comparisons of the biological reactivity of naturally occurring HMW and LMW DOM indicate that HMW DOM is generally more bioreactive (Amon and Benner, 1994,1996a). The yields of hydrolyzable neutral sugars and amino acids decrease with increasing decomposition and diagenetic alteration (Cowie and Hedges, 1994). The depth-related trends in the size, composition, and bioreactivity of DOM in the ocean indicate that most biorefractory DOM in the deep ocean exists as relatively small (1 year) periods of time. It is interesting that most of the DOM released from bacteria was not characterized at the molecular level and was therefore compositionally similar to DOM in the ocean (Ogawa et al, 2001). It is not known why some bacterially-derived DOM is resistant to biodegradation. Several recent studies have found chemical evidence for the bacterial origin of biorefractory DOM in the ocean. Membrane-derived bacterial proteins (porins) have been identified as ubiquitous components of DOM in seawater (Tanoue et al, 1995). The D-form enantiomers of common amino acids found in the bacterial cell wall polymer, peptidoglycan, have also been reported as important components of DOM in seawater (McCarthy et al, 1998). A variety of methylated sugars and amino sugars that are likely to be of bacterial origin have also been identified in marine DOM (Boon et al, 1998; Kaiser and Benner, 2000). It is difficult to estimate the fraction of DOM of bacterial origin based on these biomarkers, but there is growing evidence that a substantial fraction of biorefractory DOM in the ocean is derived from bacteria. Knowledge of the origins of biorefractory DOM will provide important clues about the mechanisms of formation of this largest reservoir of fixed carbon in the ocean.

C. FATE OF TERRIGENOUS DOM IN THE OCEAN The global annual discharge to the ocean of terrigenous DOC from rivers is quite large (0.25 Pg), yet there is no evidence for a major terrigenous component

83

84

Ronald Benner

of DOM in the ocean (Hedges et ah, 1997). This is particularly surprising because riverine DOM is generally resistant to biodegradation and has a chemical composition indicating that it is largely soil derived and highly degraded (Benner et al, 1995; Hedges et al, 1994). The stable carbon isotopic composition of DOM in the ocean indicates a predominantly marine, rather than terrestrial, origin and the low concentrations of lignin-derived phenols, unique tracers of terrestrial carbon, are also indicative of a minor terrestrial component (Meyers-Schulte and Hedges, 1986; Druffel et al, 1992; Opsahl and Benner, 1997). Concentrations of Ugnin-derived phenols that are 2.6 times higher in the Atlantic than the Pacific appear to reflect the proportionately higher riverine discharge to the Atlantic (Opsahl and Benner, 1997). In both of these ocean basins the residence time of lignin-derived phenols is relatively short (20-130 years) compared with bulk marine DOM, indicating variability in the predominant mechanisms for removal of these chemically-distinct fractions of DOM. The transport of riverine DOM to the coastal ocean is largely conservative, although losses of specific components due to flocculation can occur at low salinities during mixing (Sholkovitz, 1976; Fox, 1983; Benner and Opsahl, 2001; see Cauwet, Chapter 12). Riverine DOM is slowly degraded by microorganisms, and it appears that microbial degradation of riverine DOM continues in the coastal ocean (Chin-Leo and Benner, 1992). Riverine DOM is generally more photoreactive than marine DOM, and the combination of photochemical and microbial processes is likely to be important for its removal from the ocean (Kieber et ai, 1990; Miller and Zepp, 1995; Amon and Benner, 1996b; Miller and Moran, 1997; Opsahl and Benner, 1998). Additional field and laboratory experiments are needed to understand the mechanisms of removal of terrigenous DOM from the ocean and the factors controlling this process. Studies of the distribution of terrigenous DOM are needed to better define the geographic regions where removal processes are most active. Recent studies of the distribution and composition of terrigenous DOM in the Arctic ocean indicate that physical transport to the north Atiantic is the dominant removal mechanism from these ice-covered waters (Opsahl et al, 1999). Similar regional studies are needed in temperate and tropical environments where terrigenous DOM receives greater exposure to sunlight and warmer temperatures. Molecular evidence of photochemical transformations of terrigenous DOM in the northern Gulf of Mexico near the Mississippi River delta was recently observed (Benner and Opsahl, 2001). Improved and expanded bulk and molecular characterizations are needed to provide a more complete accounting of the components of terrigenous DOM in seawater, because the different chemical components of terrigenous DOM have varying susceptibilities to photochemical and microbial removal mechanisms.

Chemical Composition and Reactivity

85

ACKNOWLEDGMENTS I thank the U.S. National Science Foundation for supporting my research, J. Hedges and S. Wakeham for comments on the manuscript, and the students, postdoctoral associates, and technicians who have passed through my laboratory and contributed to the concepts and research presented in this chapter.

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Chapter 4

Production and Removal Processes Craig A. Carlson^ Bermuda Biological Station for Research, St. Georges GEOl, Bermuda I. Introduction II. DOM Production Processes A. Extracellular Phytoplankton Production B. Grazing-Induced DOM Production C. DOM Production via Cell Lysis D. Solubilization of Particles E. Bacterial DOM: Origination and Transformation of DOM III. DOM Removal Processes A. Biotic Consumption of DOM B. Abiotic Removal Processes IV. DOM Lability A. Biologically Refractory DOM B. Biologically Labile DOM

C. Biologically Semilabile DOM D. Continuum of Biological Lability V. DOM Accumulation A. Abiotic Formation of Biologically Recalcitrant DOM B. Biotic Formation of Recalcitrant DOM C. Limitation of Bacterial Growth and Accumulation of Biodegradable DOM D. Microbial Conmiunity Structure and DOM Utilization VI. Summary References

I. INTRODUCTION Dissolved organic matter (DOM), operationally defined as organic matter that passes a GF/F or a 0.2 /jim filter, represents one of the largest exchangeable carbon reservoirs on earth. The global dissolved organic carbon (DOC) pool is estimated to be 685 Gt C (Hansell and Carlson, 1998a), a value comparable to the mass of ^Present address: Department of Ecology, Evolution and Marine Biology, University of California at Santa Barbara, Santa Barbara, CA 93106-9610. Fax: (805) S93-4124. E-mail: carlson@lifesci. ucsb.edu. Biogeochemistry of Marine Dissolved Organic Matter Copyright 2002, Elsevier Science (USA). All rights reserved.

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inorganic C in the atmosphere (MacKenzie, 1981; Fasham et al., 2001). Net oceanic uptake of CO2 is approximately 2.1 Gt per year (Quay et al, 1992; Takahashi et al, 1999), a small percentage of the oceanic DOC pool. Small perturbations in the production or sink terms of the oceanic DOC pool could strongly impact the balance between oceanic and atmospheric CO2. Thus, processes that control DOM production, consumption, and distribution are biogeochemically significant with regard to carbon export (Copin-Montegut and Avril, 1993; Carlson et al, 1994; Ducklow et al, 1995; Hansell and Carlson, 1998b, 2001) and carbon storage in the ocean interior (Hansell and Carlson, 1998a; Hansell et al, 2001). DOM also has important ecological significance as a byproduct of biological productivity and as a substrate that supports heterotrophic bacterial growth (Williams, 1970; Pomeroy, 1974; Williams, 1981; Azam et al, 1983; Ducklow and Carlson, 1992). A number of biological processes result in the production of DOM (Fig. 1). A portion of the carbon lost from the planktonic food web as DOC is then salvaged by heterotrophic bacterioplankton, initiating the microbial loop (Azam et al, 1983), and is either repackaged into bacterial particles and passed to higher trophic levels (a trophic link), or remineralized back to its inorganic constituents (a trophic sink; Ducklow et al, 1986). While 50% of carbon fixed by phytoplankton (Ducklow and Carlson, 1992; Williams, 2000) or more (del Giorgio and Cole, 1998) is routed through DOM and processed by bacterioplankton, a fraction of DOM production accumulates in the surface waters and is resistant to rapid microbial attack. Hansell and Carlson (1998b) estimate that 1.2 Gt C year~^ or 17% of global new production escapes rapid microbial utilization, accumulates in the surface waters and is available for export to the ocean's interior. Thus, factors that control the production, removal, and accumulation of DOM in the surface ocean have both ecological and biogeochemical significance. The objectives of this chapter are to (1) review the various mechanisms of DOM production, (2) review the processes of DOM consumption and removal, (3) examine the characteristics of the various pools of DOM in terms of biological lability and their ecological and biogeochemical significance, and (4) review the factors and processes that lead to DOM accumulation. The conceptual model of the various DOM production and removal processes outlined in Fig. 1 will serve as a guide for the first two sections of this review.

11. DOM PRODUCTION PROCESSES The euphotic zone is the principal site of organic matter production in the open ocean. Net production of DOM results from the temporal and spatial uncoupling of in situ biological production and consumption processes (Fig. 1). Net DOM production is most evident in ocean regions that experience annual phytoplankton

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93

Sinking Particles & Aggregates

Figure 1 Schematic representation of the various DOM production and consumption processes in marine systems. The broken arrows associated with viruses represent viral infection and its effect on DOM production via cell lysis. The dark arrow represents dominant biotic removal process. Roman numerals and letters represent the section in the text where the process is discussed.

blooms (Table I). The magnitude and quality of DOM produced during these bloom events varies considerably and is controlled by a number of biological, chemical, and physical parameters. While DOM production is ultimately constrained by the magnitude of primary production (PP), there are several mechanisms responsible for DOM production including (a) extracellular release (ER) by phytoplankton, (b) grazer mediated release and excretion, (c) release via cell lysis (both viral and bacterial), (d) solubilization of particles, and (e) bacterial transformation and release.

Table I Comparison of Maximum Daily 14C Primary Production (PP) Rates and Seasonally Produced DOC, DON, and DOP (ADOC, ADON, and ADOP) Stocks and Concentrationsduring PhytoplanktonBlooms for Selected Marine Sites

Site

DOM deptha (m)

ADOM stock and concentration

PP (mmol m-' day-')

ADOM

type

mmolm-2

pM

548

Arctic Northeast water PolYnYa

&70

11-225

ADOC

35M76Ob

Northeast water PolYnYa

0-70

11-225

ADON

0-27b

04.39

Cornm ents ADOC, change in DOC between May and June; range represents spatial variability ADON, change in DON between May and June; range represents spatial variability; C:N of ADOM > 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.

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104

Table III Selected Examples of Conflicting Effects of Light Intensity, Nutrient Limitation, Temperature, and Community Structure on Percent Extracellular Release (PER) of DO^^C Factor Light intensity

Species Natural assemblage Phaeocystis pouchetii (culture) Natural assemblage Natural assemblage Natural assemblage Natural assemblage (diatoms)

Natural assemblage (diatoms and Phaeocystis pouchetii) Natural assemblage

Natural assemblage Chlorella vulgaris, Isochrysis galbana, Synechococcus sp. (cultures) Nutrient concentration

Natural assemblage

Chaetoceros qffinis (culmre)

Diatom dominated assemblage (mesocosm) Chaetoceros affinis (culture)

Phaeocystis sp.-dominated assemblage (mesocosm) Marine diatoms (cultures)

Effect on PER

Citation

Increased PER with high irradiance hicreased PER with high irradiance

Hellebust, 1965

Increased PER with high irradiance Increased PER with high irradiance Increased PER with high irradiance No significant difference between high and low irradiance on PER No significant difference between high and low irradiance on PER No significant difference between high and low irradiance on PER Increased PER with low light intensity Increased PER with low light intensity Increased PER with decreasing nut concentration Increased PER with decreasing nut concentration Increased PER with decreasing nut concentration Increased PER with decreasing nut concentration Increased PER with decreasing nut concentration Increased PER with decreasing nut concentration

Guillard and Hellebust, 1971 Mague et al, 1980 Wood and Valen, 1990 Thomas, 1971 Smith et al, 1977 Lancelot, 1983

Williams and Yentsch, 1976 Fogg, 1966 Zlotnik and Dubinsky, 1989 Fogg, 1966

Myklestad et al., 1989 Norrman et al. 1995 Obemoster and Hemdl, 1995 Smith et al. 1998 Goldman et al., 1992 (Continues)

105

Production and Removal Processes Table III (Continued) Factor

Community structure

Species

Citation

Natural assemblage (diatoms) Natural assemblage (diatoms) Natural assemblage

No relationship between PER and nutrient concentration No relationship between PER and nutrient concentration No relationship between PER and nutrient concentration

Smith et al, 1911 Lancelot, 1983

Phaeocystis pouchetii, diatoms Phaeocystis pouchetii, diatoms

Phaeocystis pouchetii had greater PER than diatoms No difference for PER between natural assemblages of diatoms vs Phaeocystis pouchetii No difference for PER between natural assemblages of diatoms vs Phaeocystis pouchetii

Lancelot, 1983

Phaeocystis antarctica, diatoms

Temperature

Effect on PER

Diatom culture (Leptocylidrus danicus) Chlorella vulgaris, Isochrysis galbana. Synechococcus sp. (cultures)

PER of marine diatom was temperature independent between 5 and 20° C temperatures >30°C increased ER

Williams and Yentsch, 1976

Vemet et at.. 1998

W. 0 . Smith unpublished data Verity, 1981

Zlotnik and Dubinsky, 1989

b. Passive Diffusion Model The passive diffusion model is based on the maintenance of a concentration gradient of LMW photosynthate across the autotrophic cell membrane. Monomers, such as neutral sugars, and nitrogenous compounds, like dissolved free amino acids (DFAAs), would be released via this model (Fogg, 1966). The subsequent uptake of LMW compounds by bacterioplankton and diffusion would maintain the gradient and elicit passive diffusion from phytoplankton cells (Bratbak and Thingstad, 1985; Bj0msen, 1988). Bj0msen (1988) suggested that passive exudation was a process that continued through the night and was correlated to phytoplankton biomass ("property tax") rather than photosynthetic rate. He estimated ER of carbon to be 5% of phytoplankton biomass per day. c. Model Comparison Reports in the literature support both the overflow and the passive diffusion models. For example, observations of ER at night (Mague et al, 1980; Herman and

106

Craig A. Carlson

Kaplan, 1984) and the release of low-molecular-weight compounds (Mague et al, 1980; S0ndergaard and Schierup, 198 l;M0ller-Jensen, 1983; Lee and Rhee, 1997) support the passive diffusion model. However, findings for the release of HMW extracellular products (Guillard and Hellebust, 1971; Lancelot, 1983; Lancelot and Billen, 1984; Lignell, 1990; Biddanda and Benner, 1997), absence of ER at night (Veldhuis and Admiraal, 1985), and enhanced ER during nutrient limitation (Lignell, 1990; Wood and Valen, 1990; Goldman et aL, 1992; Smith et a/., 1998) support the overflow model. In a cross-system analysis of 16 studies, Baines and Pace (1991) found that PP explained 69% of the variance of ER; a finding that supports the overflow model over ih^ passive diffusion model. It is likely that these models are not mutually exclusive and that both models are correct given the right environmental conditions and plankton community structure. The mechanisms controlling ER are still poorly understood and conflicting reports in the literature suggest that specific environmental and growth conditions may control which model is dominant at any given time.

B. GRAZING-INDUCED D O M

PRODUCTION

Jumars et al. (1989) argued that the principal pathway of DOM from phytoplankton to bacteria was via the byproducts of zooplankton ingestion and digestion. The two classes of zooplankton, metazoa (macrozooplankton) and protozoa (microzooplankton), remove significant fractions of phytoplankton production and bacterial production in marine systems. Macrozooplankton and microzooplankton remove 1-77 and 4-60% of phytoplankton production, respectively, depending on the system (Sherr and Sherr, 1988, and citations within). Caron et al. (1991) demonstrated that nanoplanktonic protists are the major consumers of prokaryotes and could remove 54 and 75% of the cyanobacteria and heterotrophic bacteria assemblage each day in field and laboratory experiments. Thus, macro- and microzooplankton grazers can potentially play an active role in transforming particulate carbon back to the dissolved phase via a variety of processes including sloppy feeding, egestion, and excretion (Table IV). The role of grazers in DOM production has been discussed previously (Ducklow and Carlson, 1992; Nagata and Kirchman, 1992; Nagata, 2000) and will be reviewed here. 1. Macrozooplankton Copping and Lorenzen (1980) found that when copepods were fed ^"^C-labeled diatoms, up to 27% of the radiocarbon ingested appeared as DOC. The four main processes by which macrozooplankton could release DOM are excretory release, egestion (release of unassimilated material), breakage of large prey during handling and feeding (sloppy feeding), and release from fecal pellets. Migrating zooplankton

Production and Removal Processes

107

in the Sargasso Sea excrete as much as 2-10% body C day~^and 1-6% of body N day"^ in the form of bulk DOC and DON, respectively (Table IV; Steinberg et ah, 2000; Steinberg et al, submitted). Nitrogen is a dominant excretory product for macrozooplankton. Inorganic nitrogen and urea are continually released by macrozooplankton (Bidigare, 1983) but DFAAs are released in pulses and make up approximately 10-21% of total N excreted (Bidigare, 1983, and citations within). Nagata (2000) speculated that the short burst in DFAA release might be more indicative of egestion than excretion but more data are needed to validate this hypothesis. Lampert (1978) showed that up to 17% of the POC removed by Daphnia pulex was lost as DOC due to cell damage during feeding. Fecal pellets may also be a source of DOM. Jumars et al. (1989) hypothesized that DOC would diffuse from fecal pellets within minutes of release. However, in experiments conducted with ^"^C-labeled fecal pellets Urban-Rich (1999) found that >50% of fresh fecal pellet carbon was released over the course of 48 h, much longer than the minutes hypothesized by Jumars et al. (1989). Lampitt et al. (1990) and Strom et al. (1997) suggested that for DOC release from fecal pellet to be significant the pellets needed to be broken upon reingestion by macrozooplankton. 2. Microzooplankton Microzooplankton grazing, consisting of herbivory and bacterivory, can release varying percentages of C, N, or P depending on prey type (Nagata, 2000). DOM released via herbivory supplies new DOM to the system, whereas bacterivory recycles DOM (Nagata and Kirchman 1992). Microzooplankton release DOM through egestion and possibly diffusion. As food vacuoles fuse with cytoplasmic walls, unassimilated DOM, undigested prey, colloidal material, and enzymes are evacuated (Nagata, 2000). This released DOM is composed of high- and low-molecularweight compounds (Taylor et al, 1985), DFAA, and dissolved combined amino acids (Nagata and Kirchman, 1991) as well as dissolved organic phosphorus (DOP; Caron et al, 1985; Andersen et al., 1986). The production of DOM by microzooplankton is positively correlated with the food concentration, and its release is highest during exponential growth (Nagata and Kirchman, 1992). Hagstrom et al. (1988) hypothesized that DOM release by microzooplankton was the dominant pathway of C flow to bacterioplankton. Nagata (2000) hypothesized that if or when phytoplankton growth is balanced by protozoan grazing, then 10-30% of particulate PP could be transformed to DOC with bacterivory being an additional source of DOM. 3. Biogeochemical Significance Zooplankton grazing has biogeochemical as well as ecological significance. For example, bacterivory increases nutrient regeneration efficiency within the

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WilUams, 1995 Daley era/., 1999 Ogawaera/., 1999 Carlsonef al, 2000 Hansen and Carlson, 2001

Turnover on time scales of centuries to millennia; representative of deep water concentrations

WiUiams and Druffel, 1987 Bauer era/., 1992

Dominated by LMW compounds

Bennerera/., 1992 Amon and Benner, 1996

Resistant to microbial remineralization

Barber, 1968 Ogura, 1972

Most diagenetically altered; low carbohydrate and aldose yield

Cowie and Hedges, 1994 Skoog and Benner, 1997

C:N (molar ratio) 14-20

Ogawaera/., 1999 Carlson et ai, 2000 Hansen and Carlson, 2001

Mean age 4000 years

Bauer era/., 1992

Photochemically active material that can be transformed to biologically labile material

Kieber era/., 1989 Mopper era/., 1991 Moran and Zepp, 1997

Observed gradient in deep ocean DOC between North Atlantic and North Pacific (i.e. 14 fxM C)

Hansen and Carlson, 1998a

Most refractory portion of DOC is composed of approximately 34/LtMC

Hansen and Carlson, 1998a

Portion of deep DOC that turns over on time scale of ocean mixing; varies with location

Hansen and Carlson, 1998a

into the lability of DOM found in the ocean. See Benner (Chapter 3) for extensive review of DOM composition and reactivity.

A. BIOLOGICALLY REFRACTORY DOM The refractory pool is the largest fraction of the bulk DOC. It is dominated by diagenetically altered LMW DOM (Benner et ai, 1992; Amon and Benner, 1996; Skoog and Benner, 1997). This finding would appear contradictory because

Production and Removal Processes

127

bacterioplankton can take up only LMW compounds directly; however, not all LMW compounds are bioavailable. Amon and Benner (1996) proposed a sizereactivity continuum model which suggests that as organic matter is decomposed it become less bioreactive and smaller in size; thus bioreactivity of organic matter decreases as follows: POM - ^ HMW DOM —> LMW DOM, where each size fraction consists of a continuum of composition, reactivities, and diagenetic states (Amon and Benner, 1996). The refractory pool is best represented by the deep ocean DOC stocks (>1000 m), with an apparent mean age of 4000 to 6000 years old in the North Atlantic and the North Central Pacific oceans, respectively (Williams and Druffel, 1987; Bauer et al, 1992). Because the mean age of the deep DOC is much greater than the time scale of thermohaline circulation refractory DOC is reintroduced to the surface waters as it follows the path of ocean circulation. Based on mass balance calculations and natural A^^C estimates, Druffel et al. (1992) suggested that the most refractory component of the bulk DOM pool was uniformly distributed throughout the water column and represents approximately 70% of surface DOC in thermally stratified systems (Carlson and Ducklow, 1995, 1996; Cherrier etal, 1996). In polar systems, the surface contribution of refractory DOM can be even higher due to the absence or reduced concentration of "semilabile" DOM (see below and Hansell, Chapter 15). Although this material is biologically resistant there must be a removal mechanism or DOC would continually accumulate and the apparent age would be much older. While deep DOC is biologically refractory (Barber, 1968), the deep ocean DOC gradient of 14 fjM from the North Atlantic to the North Pacific (Hansell and Carlson, 1998a) indicates that a portion of the deep ocean DOC is removed on time scales of ocean mixing. Rate constants required to account for continuous removal are on the order of 10~^ to 10""* year~\ too slow for continuous microbial decomposition to be likely (Anderson and Wilhams, 1999; WilHams, 2000). Anderson and Williams (1999) proposed a coupled biological-photochemical model in which deep DOC is biologically refractory but photochemically reactive. Once exposed to surface UV irradiation, it is removed via photooxidation or broken into labile compounds and removed via microbial remineralization (Kieber etal, 1989; Mopper et al, 1991). This photochemical mechanism most likely accounts for a portion of refractory DOM loss (Moran and Zepp, 1997); however, as Williams (2000) points out, this process is restricted to the surface ocean and cannot account for the deep DOC gradient observed by Hansell and Carlson (1998a). Williams (2000) proposed a third mode of deep DOC decomposition, in which attached bacteria associated with sinking particles generate short bursts of microbially mediated DOM removal. Such remineralization of refractory DOM by attached bacteria, together with sorption of HMW DOM to sinking particles (Druffel et al, 1996,

128

Craig A. Carlson

1998), may help explain removal of deep refractory DOM as it moves through ocean basins via thermohaline circulation. However, Turley (1993) found that by increasing pressure to 200 atm, incorporation of [^H]leucine and [^H]thymidine (indices of bacterial productivity) was reduced by 94 and 70%, respectively, relative to incorporation rates measured at 1 atm. Based on this experiment, Turley concluded that it was unlikely that bacteria, originating from the surface waters and attached to rapidly sinking particles, play a role in particle remineralization below 1000-2000 m. Thus, further research is needed to fully elucidate refractory DOM removal mechanisms and associated rates.

B. BIOLOGICALLY LABILE DOM The most biologically reactive organic components in seawater include dissolved free compounds such as neutral monosaccharides and DFAAs. Rapid turnover (minutes to hours) maintains these compounds at nanomolar concentrations in the open ocean (Fuhrman and Ferguson, 1986; Rich et al, 1997; Keil and Kirchman, 1999; Skoog et al., 1999). Despite low concentrations, rapid turnover rates suggest that the fluxes of these compounds can be high (Fuhrman and Ferguson, 1986; Rich et ai, 1996; Keil and Kirchman, 1999; Kirchman et al, in press). For example, DFAA uptake has been shown to support 4-41% of bacterial N demand in open ocean environments and up to 100% in coastal studies (Kirchman, 2000). Glucose uptake represented 27-35% of net bacterial production in the equatorial Pacific (Rich et al, 1996) and up to 100% in the Arctic (Rich et al, 1997). In the Gulf of Mexico and the Antarctic, glucose uptake supports < 10% of bacterial production (Skoog et al, 1999; Kirchman et al, in press). Less than 5% of BCD in the Sargasso Sea (Keil and Kirchman, 1999) is supported by glucose. These latter studies indicate that compounds of slower reactivity, such as proteins and other carbohydrates (i.e., polysaccharides), support a significant portion of heterotrophic BCD with turnover on time scales of days. Seawater culture experiments, combined with direct measurements of DOM consumption and bacterial production, are often used to assess the availability of naturally occurring substrates to bacterioplankton (Fig 6; Barber, 1968; Ogura, 1972;Kroer, 1993; Carlson and Ducklow, 1996;Cherrier^^fl/., 1996;Kahler^ra/., 1997; Carlson et al, 1999). The proportion of labile DOM varies spatially with an apparent gradient of decreasing concentrations of labile DOM from coastal to oceanic systems (S0ndergaard and Middelboe, 1995). In some oceanic systems DOM production and consumption processes are so tightly coupled that there is no measurable surplus of labile DOM (on the /xM C scale) available to surface water microbes on time scales of days (Fig. 6b; Carlson and Ducklow, 1996; Cherrier et al, 1996; Carlson et al, in press).

Production and Removal Processes

129

Figure 6 Bacterial carbon production (closed circles) and DOC utilization (open circles) monitored in two seawater culture experiments conducted in the northwestern Sargasso Sea. (a) At the time this experiment was conducted, in situ DOC production and consumption processes were uncoupled, resulting in accumulation of labile DOC above the average mixed layer DOC concentration of 69 /xM C. In the presence of surplus labile DOC, bacterioplankton were able to grow rapidly and consume DOC at an efficiency of approximately 14%. (b) At the time this experiment was conducted, in situ production and consumption processes were tightly coupled preventing accumulation of labile DOC (on the /xM scale). The limited availability of labile DOC at the time water was collected for this experiment prevented rapid microbial growth and DOC degradation. Figure adapted with permission from Carlson and Ducklow (1996).

C

BIOLOGICALLY SEMILABILE D O M

DOM stocks in excess of the deep refractory pool are composed of "labile" and "semilabile" DOM. Because labile DOM concentrations represent a very small fraction of bulk DOM in the open oceans (0-6%), the vertical gradient of the bulk DOM observed in thermally stratified systems is mostly comprised of semilabile DOM (Fig. 5; Carlson and Ducklow, 1995; Cherrier et al, 1996). The semilabile fraction of DOM is biologically reactive only over months to years, allowing it to accumulate in the surface waters. It is this DOM fraction that can be an important export term provided it escapes microbial degradation in the surface waters long enough to be entrained to depth via convective mixing (Copin-Montegut and Avril, 1993; Carlson et al, 1994; Hansell and Carlson, 2001) or advection along isopycnal surfaces (Hansell etal, 2001). For example, in the Sargasso Sea overturn of the water column serves to export DOC that escapes rapid microbial degradation to depths greater than 100 m, with the amount exported dependent on the depth of mixing. Time-series measurements of DOC conducted from 1992-1998 indicated that DOC export ranged from 0.4 to 1.4 mol C m"^ year~^ (Carlson et al, 1994; Hansell and Carlson, 2001). In some years the contribution of organic matter exported out of the euphotic zone in the dissolved phase was greater than that export as POC, as measured by sediment traps (Carlson etal, 1994).

130

Craig A. Carlson

Semilabile DOM is composed of both LMW and HMW compounds of varying lability. HMW compounds can be biologically reactive (Amon and Benner, 1994, 1996), but must be hydrolyzed to monomers via extracellular enzymes prior to bacterial uptake. The time required for microbes to synthesize and release these enzymes decouples DOM production and bacterial remineralization processes (Billen and Fontigny, 1987), resulting in transient DOM accumulation. Dissolved carbohydrates totaling as much as 40% of the surface DOC stocks comprise the largest identifiable fraction of the bulk DOC pool (Bumey et al, 1979; Benner et ai, 1992; Pakulski and Benner, 1994; B0rsheim et al, 1999). Nitrogenous compounds such as amino acids account for a relatively minor component of DOM in the ocean (McCarthy et ai, 1996). The major classes of carbohydrates present in seawater include amino sugars, uronic acids, and neutral sugars (Borch and Kirchman, 1997; Skoog and Benner, 1997; Kirchman et al, 2001, and citations within). Polysaccharides comprise a major fraction of reactive HMW DOM in oceanic surface waters (Benner et ai, 1992; McCarthy et al, 1996) and decrease in concentration with depth.

D. CONTINUUM OF BIOLOGICAL LABILITY In thermally stratified oceanic sites there exists a broad continuum of lability for the semilabile DOM pool, with some fraction of semilabile DOC turning over on seasonal time scales and other fractions persisting for years (Anderson and Williams, 1999; Carlson etal, 2000). Time-series measurements of DOM dynamics can provide insight into turnover rates of various DOM pools. Turnover rates of semilabile DOC for specific depth horizons can be defined as the difference between the maximum and minimum DOM stocks within that depth horizon over an annual cycle. For example, in the Sargasso Sea approximately 4.8 mol m~^ (30%) of bulk DOC measured in the surface 250 m (maximum mixed layer depth for time period) is composed of semilabile DOC yet a maximum of only ^ 1 mol m~^ (20% of semilabile DOC) turns over on a time scale of 1 year or less (Fig. 7; Carlson et al, 1994, 2000; Hansell and Carlson, 2001). The percentage contribution of reactive compounds such as carbohydrates decreases with depth (Benner et al, 1992; Pakulski and Benner, 1994; Skoog and Benner, 1997), indicating that deeper semilabile DOM is more recalcitrant than that found in the surface waters. Because a portion of the semilabile DOM pool present at depth is transported via isopycnal mixing, the ventilation age of water provides insight as to the turnover time of the deeper semilabile DOM. Ventilation ages of water between 250 and 500 m in the Sargasso Sea can range from 1 to > 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

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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

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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

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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

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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

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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).

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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.

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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.

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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.

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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.

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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

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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 (

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Latitude (deg N) Figure 2 (Top) Concentrations of dissolved inorganic nitrogen (DIN) and total organic N (TON) and (Bottom) the total organic carbon (TOC) to TON ratio along a transect in the North Pacific (modified with permission from Abell et al, 2000).

et al, 1997), the oligotrophic North Pacific (Malta and Yanada, 1990; Harrison et al, 1992; Loh and Bauer, 2000), along a transect from Bermuda to Chesapeake Bay (Bates and Hansell, 1999), the Arctic (Wheeler et al, 1997), the North Pacific (Abell etal, 2000), and the equatorial Atlantic (Vidal etal, 1999). In the Southern California Bight, DON concentrations are generally uniform in the surface layer but become more variable and often increase within the nitracline (Hansell et al, 1993). Elevated DON concentrations at the surface suggest that DON can be exported to depth (Toggweiler, 1989; Hopkinson et al, 1997). DON export, and subsequent ammonification and nitrification, is estimated to supply 19% of remineralized NO3" at depth off Georges Bank (Hopkinson et al, 1997); 15% in the Bering Sea

172

Deborah A. Bronk

(Koike and Tupas, 1993), 10% at various sites in the Pacific (Jackson and Williams, 1985), and 25% in the North Pacific (Malta and Yanada, 1990). Vidal et al (1999) calculated vertical gradient-driven fluxes of DON in the equatorial Atlantic using vertical profiles of DON concentrations and estimates of a vertical eddy diffusion coefficient. They found that surface DON did appear to be transported to depth at times, however, the direction of the dominant flux varied along the north-south transect. The DOC pool is commonly envisioned as being composed of a labile, semilabile, and refractory component based on vertical profile data (Kirchman et al, 1993; Carlson and Ducklow, 1995; see Carlson, Chapter 4). In a similar fashion, the refractory TON pool is estimated at 4 /xM in equatorial Pacific waters, based on TON concentrations in the deep ocean (Libby and Wheeler, 1997). Assuming this recalcitrant fraction is present throughout the water column, refractory TON is -- 60% of the TON in the upper 40 m; the C:N of the refractory TON is 9.9. The semilabile pool in the upper 40 m ranges from 3.4 to 5.8 /xM TON and has a lower C:N ratio (5.1 to 8.5) than the refractory pool. 4. Seasonal Variations The question of whether DON concentrations exhibit a seasonal pattern is an open one. There are areas where no seasonal pattern is indicated, including the Santa Monica Basin (Hansell et al, 1993) and at the Bermuda Atlantic Time Series (BATS) site in the Sargasso Sea (Hansell and Carlson, 2001). In contrast, some studies suggest that DON increases in late spring and summer, including work in the Gulf of Mexico (Lopez-Veneroni and Cifuentes, 1994), Chesapeake Bay (Bronk et al, 1998), North Inlet, SC (Lewitus et al 2000), and a suite of rivers draining into the Baltic Sea (Stepanauskas et al, in press). The most compelling evidence for a seasonal cycle is presented by Butler et al (1979), who conducted an 11-year study of DON concentrations in the English Channel. They documented a steady increase in DON concentrations from January through August and then a steady decline from August to December. The clear seasonal pattern observed by Butler et al (1979) highlights the importance of long-term data sets in defining these types of patterns. In general, oligotrophic environments do not show seasonality in DON, as would be anticipated given the scant supply of N. 5. Link between DON Distributions and Inorganic N There have been a number of observations linking decreases in NOs" or elevations in N2 fixation and accumulations of DON in near surface waters. In the Pacific, an increase in DON/TON concentrations and a concomitant decrease in NOs" concentrations as one moves away from the equator is shown by a number of studies (Libby and Wheeler, 1997; Hansell and Waterhouse, 1997; Raimbault

Dynamics of DON

173

et al, 1999). This pattern suggests that new TON production is fueled by high equatorial new production and subsequent organic N release. An estimated 37 ± 14 and 81 ± 54% of net NO3 ~ depletion accumulates as DON during the movement of upwelled equatorial water to the north and south, respectively (Libby and Wheeler, 1997). In the Ross Sea, approximately 10% of the net NO3" draw down in surface waters accumulates as DON (Carlson et al, 2000). Similar relationships between NOs" consumption or disappearance and TON/DON production have been observed in the English Channel (Butler et al, 1979), in the subarctic Pacific (Malta and Yanada, 1990), in Chesapeake Bay (Bronk et al, 1998), and in a coastal pond {Co\\o% etal, 1996). In the Mississippi River plume, Benner et al (1992a) used 5^^N data to demonstrate the conversion of NO3" to high-molecular-weight (HMW) DON, isolated using a > 1-kDa ultrafiltration unit. The DON produced is likely the result of phytoplankton uptake and subsequent conversion to DON that is then released. The 5^^N of the DOM pool is ~3%o at both the river and Gulf endpoints. At intermediate salinities, however, the 5^^N of the DOM pool increases to ^9%o. The likely cause for this increase is the conversion of NOs" to DON; the NOs" in this region has a ^^^N of -^10%^. In the Pacific, Atlantic, and Gulf of Mexico, the ^^^N of the HMW DON ranges from a mean of 7.9 ± 0.7 in the Pacific to 9.9 ± 0.5%o in the Gulf of Mexico (Benner et al, 1997). There is no relationship between the concentration of HMW DON and the 5 ^^N of the material, suggesting that the isotopic signature reflects the source of the N, not isotopic fractionation during decomposition. The lowest (5^^N values (6.6%o) are observed in the surface waters at the BATS site in the Sargasso Sea, and are likely the result of the use of isotopically light new NOs", which has a 5^^N of ~3.5%o in the region (Altabet, 1988). These low values can also indicate the addition of isotopically depleted N from N2fixation(atmospheric N2 is ~0%o) and then subsequent release of DON from N2fixerssuch as Trichodesmium (see Section III. A.2). In the Pacific, the lightest 5^^N values are measured at the equator where new NOs" is upwelled (Benner, et al, 1997). The 5^^N of DON increases to the north and south of the equator, suggesting that the DOM is being produced biologically during meridonal transport as described above. In the reverse of the NOs" to DON conversion discussed above, Kemer and Spitzy (2001) documented that between 75 and 100% of the LMW DON and NH4+ is consumed and converted to NOs" in the Elbe estuary via nitrification. This is the analogous process that produces the inverse relationship between DON and NOs" concentrations when moving from the surface into the deep ocean (Lara etal, 1993). Finally, there are examples where increases in DON concentrations do not correlate well with decreases in NOs". For example, along the NW African coast, Vidal et al (1999) found that the downward flux of DON exceeds the upward supply of N03~, indicating an additional supply of N from meridonal transport

174

Deborah A. Bronk

(Libby and Wheeler, 1997) or perhaps atmospheric inputs (Cornell et al, 1995). At the station with the highest DON flux, in excess of NOs" influx, there were significant concentrations of Trichodesmium. C. CHEMICAL COMPOSITION OF THE DON

POOL

DON is a heterogeneous mixture of compounds composed of biologically labile moieties, which likely turn over on the order of days to weeks, and refractory components, which persist for months to hundreds of years, and comprise the bulk of the DON measured in the ambient DON pool (see Section V). The more refractory forms are quantitatively dominant with respect to ambient concentrations, but their importance as a potential N source is far exceeded by the smaller, more labile compounds. A large number of compounds have been identified within the DON pool, including urea, dissolved combined amino acids (DCAAs), dissolved free amino acids (DFAA), humic and fulvic substances, and nucleic acids. The remainder of the DON pool is a heterogeneous mixture of unidentified compounds. In this section, individual organic compounds are discussed including measurement techniques and a review of recently measured concentrations, followed by a discussion of recent efforts to chemically characterize the HMW fraction and research priorities for the future. 1. Urea Urea is a low-molecular-weight (LMW) organic compound, which is a product of organic matter decomposition and organismal excretion. The two methods commonly used to measure urea concentrations are the urease method, which involves enzymatic hydrolysis of urea to carbon dioxide and ammonia (McCarthy, 1970), and the direct colorimetric measurement of urea using diacetyl monoxime (Price and Harrison, 1987). Note that urea is occasionally treated as an inorganic N form (Capone, 2000; J0rgensen et al, 1999). The rationale for this practice is that urea is used primarily as a N source and not as a form of energy. Out of solidarity with poor researchers assembling large tables, I suggest this practice be discontinued, because it confuses the calculation of DON concentrations. Urea contains C, as well as N, and so should sit squarely in the DON pool. Urea concentrations range widely from 0 to 13 /xM in the studies surveyed (Table II). In general, urea concentrations in open ocean systems tend to be very low (

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B. METHODS FOR ESTIMATING BIOTIC DON UPTAKE DON is difficult to study as a N source because it is composed of a large number of compounds and the exact composition is unknown (Gardner and Stephens, 1978; Sharp, 1983; Antia et al, 1991). As a result, measurements of DON uptake rates have largely been limited to a few compounds which have commercially available ^^N, ^"^C, or ^H tracers such as amino acids or urea (Fuhrman, 1987; Hansen and Goering, 1989; Wheeler and Kirchman, 1986; Cochlan and Harrison, 1991; Antia et al, 1991). Bronk and Gilbert (1993a) developed a method for manufacturing ^^N-labeled DON produced in situ, which involves incubating a whole water sample with ^^N-labeled NH4+or NO3". The recently released DO^^N

212

Deborah A. Bronk

is then isolated using ion retardation resin (see Section III.B.2) and subsequently used as a tracer to quantify DON uptake rates. Much of the recent work focusing on the bulk DON pool used a bioassay approach, where changes in ambient and added DON are monitored over time (Seitzinger and Sanders, 1997). Bioassay approaches are particularly useful in determining the biological availability of more recalcitrant organic N compounds, such as humic substances (Carlsson and Graneli, 1993; Carlsson et ai, 1995).

C. LITERATURE VALUES OF DON UPTAKE IN A Q U A T I C E N V I R O N M E N T S

Consideration of organic N uptake is slowly becoming a routine part of many field programs. Here, uptake rates of bulk DON (i.e., the total DON pool), urea, DCAA and DFAA, humic substances, and other DON compounds are discussed. 1. Bulk DON Most of the work done on bulk DON utilization has been in freshwater systems using a bioassay approach. This work suggests that 12 to 72% of the DON pool is bioavailable on the order of days to weeks. In the Delaware and Hudson Rivers, 40-72% of the DON is consumed during 10- to 15-day dark bioassays, and DON consumption results in both an increase in PN and the release of DIN (Seitzinger and Sanders, 1997). These data suggest that the bioavailable DON can be utilized within estuaries with residence times on the order of weeks to months. In systems where residence times are shorter, riverine DON will be a source of bioavailable N to coastal waters. Stepanauskas et ai (1999a) measured the concentration and bioavailability of three MW DOM fractions in samples collected seasonally in Swedish wetlands. The percentage of bulk DON represented by the different fractions ranges from a mean of 23 % for the HMW fraction to a low of 6% for LMW DON. They found that bioavailable DON is higher in seawater than in freshwater and that bioavailability does not correlate with the C:N ratio of the DOM. The percentage of the different fractions that are bioavailable in seawater cultures are 12 ± 4, 7 ± 3, 5 ± 4, and 16 ± 8% for bulk, HMW, intermediate, and LMW DON, respectively. In additional studies in wetlands, the addition of natural DON stimulates cell-specific AMPase activity; refractory and humic-rich DOM causes a stronger stimulation than other forms believed to be more labile (Stepanauskas et al, 1999b). AMPase activity is twofold higher in seawater, relative to freshwater, indicating that hydrolysis and turnover of terrestrial DON may increase when it enters the coastal ocean (Stepanauskas ^r (2/., 1999b).

Dynamics of DON

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In two streams in Sweden, 19-55% of the bulk DON is bioavailable in short-term bioassays (Stepanauskas et ai, 2000). Only 5-18% of the DON is identified as urea, DCAA, or DFAA, suggesting that bacteria also utilize other organic N compounds. Potential DON bioavailability is positively correlated with the concentration of DCAA and the proportion of L-enantiomers in amino acids. In 7- to 8-day bioassay experiments in the Gulf of Riga, an average of 77% (8-136%) of the bacterial N biomass accumulation is a result of DCAA and DFAA uptake, and 13% of the DON is bioavailable during the study (J0rgensen et al, 1999). Bronk and Gilbert (1993a) used ^^N-labeled DON produced in situ in Chesapeake Bay and found that during the decline of the spring bloom, uptake rates of DON are higher than uptake rates of NH4+ and NO3". In August, rates of DON uptake are again higher than uptake rates of NOa", though not higher than NH4"^. 2. Urea In general, phytoplankton are believed to be the primary users of urea in marine systems (Price and Harrison, 1988, Table VI). More recent studies, however, have called this belief into question (Tamminen and Irmisch, 1996). In the Thames Estuary, the addition of a broad procaryotic inhibitor reduces dark uptake rates of amino acids by 49 di 20% and urea by 86 di 25%, suggesting that, contrary to popular belief, autotrophs use a significant fraction of the amino acids and that bacterial uptake of urea is substantial (Middelburg and Niewenhuize, 2000). The whole water microbial conmiunity and the heterotrophic bacterial community alone appear to prefer amino acids, with NH4+ and urea next, and NOs" as the least preferred N substrate. In the bioassay study described in the previous section, J0rgensen et al (1999) also found that urea uptake by bacteria can be as important as DFAA uptake. In the Chesapeake Bay plume, urea contributes 60 to 80% of the N uptake measured throughout most of the year (Gilbert et al, 1991). Lomas et al (in press) reviewed urea uptake rates for over a decade in Chesapeake Bay and found that urea is consistently an important N source for the plankton community, and that the highest mean baywide rates are observed during the sunmier. Illustrating the close coupling between urea uptake and urea regeneration, Hansen and Goering (1989) found that urea uptake rates based on urea disappearance are an average of 140% greater than those based on rates of N accumulation in the Bering Sea. Because urea regeneration is prevalent in their samples, correcting for isotope dilution increases measured uptake rates by an average of 54%. 3. DCAA and DFAA Bacteria are generally considered the primary users of DCAA and DFAA. As noted for urea above, changes in the size of the DFAA pool are generally small even when rates of uptake and release are substantial, indicating that uptake and release

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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.

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232

Deborah A Bronk

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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.

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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.

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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

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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 (200 m (i.e., the pelagic marine environment which is the stated focus of this review). This otherwise unedited Global Open Ocean DOP (or GOOD) database, which includes n = 139,747 measurement pairs is available along with our enhanced DOP summary upon request of the senior author. The GOOD database includes measurements from all major ocean basins but has several large data gaps, most notably the Eastern North Pacific Ocean, the South Pacific Ocean, and the Southern Ocean (Fig. 7). Nevertheless, this n= 139,747 pairs of SRP/TDP measurements greatly exceeded our initial expectation, especially considering the relatively few open ocean DOP profiles that have been published in the refereed literature. A notable exception, that we highlight here, is the extensive survey of the North and South Atlantic Ocean basins conducted as part of the International Geophysical Year (IGY). During this 2-year (1957-1958) investigation, McGill (1963, 1964) compiled what amounts to the most comprehensive study of oceanic DOP yet attempted. This must be considered the exception to the otherwise sparse open ocean database.

A. REGIONAL AND DEPTH VARIATIONS IN DOP There are several general features of the global ocean DOP distributions. First, concentration versus depth profiles in the open ocean almost exclusively reveal

282

Karl and Bjorkman SRP (^iM)

DOP (^iM)

DOP (% TDP)

20

40

60

80

100

Figure 8 SRP, DOP and DOP as a percent of TDP for the Hawaii Ocean Time-series Sta. ALOHA (22°45'N, 158°W). These data are mean values and 95% confidence intervals based on samples collected during 110 research cruises between October 1988 and December 1999.

elevated DOP in the upper 0-100 m of the water column. For example, the 12-year climatology of SRP and DOP at Station ALOHA (22°45'N, 158°W) documents a characteristic inverse depth relationship with high concentrations of DOP in the surface water, decreasing with water depth, and vice versa for SRP concentrations (Fig. 8). The only possible exception to this general pattern might be for highlatitude habitats in winter where deep vertical mixing and low solar irradiance preclude contemporaneous near-surface DOP production via primary production. Despite this predictable depth dependence of total DOP in the open sea, individual DOP compounds or compound classes can have one of three fundamentally different distributions as a function of depth in the water column: (1) local enrichments near the sea surface with decreasing concentrations beneath the euphotic zone (similar to total DOP), (2) near surface depletion with increasing concentrations beneath the euphotic zone, or (3) constant concentration with depth. Dissolved

283

Dynamics ofDOP

nucleotides (e.g., ATP) conform to the first pattern of depth distribution and dissolved vitamins (e.g., vitamin B12) generally conform to the second pattern. These depth distributions are a result of net production/consumption and import/export processes (see Section VIII). At the subtropical location of Sta. ALOHA, DOP in the upper 100 m of the water colunm averages 0.20-0.22 /xM or approximately 70-80% of the TDP pool. Below 100 m, DOP decreases with increasing water depth to values 300 m are consistently < 10% of the corresponding TDP, indicating a deep water dominance by SRP (Fig. 8). Similar patterns are also observed for both the IGY (North and South Atlantic Oceans) and GOOD data sets (Fig. 9), with the exception of a generally lower percentage of DOP in near surface

DOP (% TDP)

DOP (% TDP) 0

1

1

•

^

1

100

200

300

• • • •

-

400

500 Q. •

Q 600

-

• '•

700

-

• •

.

\

* •

IGY

800 -

900

nnn

•

WORLD OCEAN

- • •

1

1

1

1

Figure 9 Vertical profiles of DOP, expressed as percentage of TDP, for the entire Global Open Ocean DOP (GOOD) database and International Geophysical Year (IGY) subsampled data (see text for more details). The data are presented as mean values and 95% confidence intervals. The data were depth binned, as shown, prior to the determinations of the sunmiary statistics. For the IGY profile, n ranged from 32 (125 m) to 594 (surface) with a median ofn= 130. For the GOOD profile, n ranged from 805 (850 m) to 12,234 (25 m) with a median of n = 2700.

284

Karl and Bjorkman

waters (i.e., 50-60% of the TDP pool compared to 70-80% for the subtropical North Pacific; Fig. 8). For all oceanic DOP profiles, however, both the absolute DOP concentration and the DOP/TDP ratio vary systematically, and therefore predictably, with water depth. It is well known that the concentration of SRP, especially in subthermocline (>600 m) waters, increases as the age of the water mass also increases (Levitus etai, 1993). This is a result of the long-term net remineralization of organic matter. These spatial patterns are clearly evident in the IGY data set (Figs. 10 and 11 [see color plate]). A comparison of two zonal sections, one at 40°N and the other at 24.25°S in the Atlantic Ocean documents the following: (1) a contrast between relatively low TDP/low SRP waters at 40°N compared to high TDP/high SRP in the northward flowing Antarctic Bottom Water in the west (>4000 m) and Antarctic Intermediate Water (^1000 m) seen at 24.25°S; (2) nearly homogeneous concentrations of all forms of P in the relatively "young, well mixed" North Atlantic compared to the South Atlantic, especially for the subeuphotic zone waters; and (3) a general increase in surface water DOP concentrations in the South Atlantic basin, especially in the near-surface water (Figs. 10 and 11). For DOP there also appears to be a significant basin-scale east-to-west gradient, at least at 40°N, with higher DOP concentrations in the western North Atlantic, and minimum DOP concentrations in the deep central waters of both basins. These regional variations are superimposed on the general global DOP distributions described previously and are probably related to circulation patterns and processes. A similar Atlantic Ocean intrabasin gradient in DOP was recently reported by Vidal et ah (1999) for a transect from the Canary Islands to Argentina (22°N to 3 PS). Euphotic zone (0-200 m) DOP concentrations in the western portion of the basin were approximately two to three times greater than they were in the eastern portion of the Atlantic basin (0.2 to 0.3 MM versus 0.25 /xM) in most coastal and estuarine habitats that have been investigated (Table II). In the semienclosed Baltic Sea and surrounding coastal regions (Fig. 13 [see color plate]), surface DOP concentrations vary considerably from a minimum, "background" concentration of 0.4 to 0.6 /xM to values >1 /xM in regions that are likely impacted by point source and non-point-source nutrient inputs. In particular, the west coast of the Jutland Peninsula appears to be especially enriched in DOP relative to offshore regions. Even in the open ocean pelagic ecosystem off the west coast of South America, elevated near surface DOP concentrations are apparent (Fig. 14 [see color plate]). Upwelling-induced organic matter production and coupled DOP production, combined with coastal runoff and locally restricted flushing can all contribute to both local and regional elevations in surface DOP.

B. D O P

CONCENTRATIONS IN THE DEEP SEA

The precision of DOP estimation is eroded when SRP is >90% of the TDP pool, for example at all ocean depths greater than approximately 500 m in the global open ocean (Fig. 9), as well as in many high-latitude surface waters. Small relative errors in SRP and TDP determinations translate into large errors in the calculation of DOP. Ketchum etal. (1955) presented a comprehensive assessment of the analytical and statistical considerations for samples collected in the equatorial Atlantic Ocean. From a paired Pi and TDP (using the method of Harvey, 1948) measurement suite that consisted of more than 1000 seawater samples, they concluded that unless the difference (i.e., TDP-P/) exceeded 10% of the TDP value, the DOP (technically, SNP) estimate cannot be considered to be significantly different from zero. For their analyses, 95% of the surface water samples contained significant DOP decreasing through the region of the phosphate-cline where P/ increases and DOP decreases with increasing water depth. At depths greater than 1000 m there was no measurable (statistically significant) DOP present; only 13 out of 259 deep water samples (5%) gave positive differences that exceeded 1 standard deviation

Dynamics ofDOP

287

(Ketchum et al, 1955). It could not be determined whether DOP was present at concentrations below the analytical detection limit, or whether DOP was truly absent. Consequently, few reUable data sets exist for DOP concentrations in the deep mesopelagic and abyssopelagic zones (>1000 m). This is quite unfortunate because the poorly understood, stepwise conversion of PP and DOP to P/ is a key metabolic process in these regions. This analytical uncertainty, for better or for worse, has not prevented ocean researchers from sampling, analyzing, and reporting subeuphotic zone DOP concentrations. The caution we urge here is to be cognizant of the statistical constraints on the methodologies employed, as they will clearly affect the ecological implications of the data obtained. Examination of these data sets documents a fairly broad range in mesopelagic and deep sea DOP concentrations that cannot be easily reconciled with any known oceanic processes. A difference of just 20-50 nM DOP between these determinations, when scaled to dimensions of the deep sea, creates or ehminates a DOP pool that becomes significant for global ocean P budgets. It is imperative that we obtain reliable deep water DOP measurements if we ever hope to understand subeuphotic zone DOP dynamics or marine P cycle as a whole. In theory, one might anticipate a small, butfiniteDOP pool that would represent a balance between local DOP production and utilization processes. The supply of DOP to depth depends to a large extent on the nature of organic matter export processes (e.g., downward diffusion and mixing of DOM versus gravitational settling of POM) and on the pathways and coupling between export and remineralization mechanisms. However, only the process of gravitational settling of particulate matter can export significant amounts of "fresh" organic materials to subthermocline (>500 m) ocean depths. During the 1- to 2-month-long journey to the seabed in the open ocean, these exported particles are disaggregated, hydrolyzed, and otherwise reworked with a continuous, and sometimes variable, loss of organic C, N, and P. Much of this organic matter is remineralized at depth which accounts for the generally increasing concentrations of dissolved inorganic nutrients (see Fig. 8 and 12). For open ocean habitats, the flux of organic P from the euphotic zone (150 m reference depth) is approximately 5 mmol P m~^ year"^ This statistical population of sinking particles is attenuated nearly an order of magnitude by the 1000-m depth horizon and nearly two orders of magnitude at the seabed (5000 m). This attrition of organic P from the sinking particulate pool as a predictable function of depth is the primary starting material for the suspended particulate and dissolved organic matter pools beneath the euphotic zone. Only when a sufficient number of determinations are available can the statistical significance of DOP in the deep sea be assured. Repeated observations of SRP and TDP in a section between Montauk Point, New York, and Bermuda during the period 1958-1961 have demonstrated the appearance, at depth, of a low-salinity subarctic intermediate water mass that covaries with DOP concentration (McGill et al., 1964). For water samples collected

288

Karl and Bjorkman Table II

Selected Marine Dissolved Organic Phosphorus (DOP) Concentrations and DOP as Percentage of Total Dissolved Phosphorus (% of TDP) Reported from a Variety of Geographic Locations Depth (m)

Method^

DOP^ (/xM)

0-1000 0-200

UV UV

0.11-0.29 0.05-0.90

Surface

DA-AH

Surface 10 Surface

PO PO

0.37 ± 0.02 Range 0.04-2.03 («=183) 0.224-0.264 0.318-0.337 0.12-0.17

Surface

PO-AA

0.2-O.6

3 11 16

PO

0.48 ± 0.01 0.63 ± 0.03 0.98 ± 0.05

5 20

UV

0.21 0.22

0-50 75-1300

UV

0.27 to 0.42 -0.15

10 25 75

PO

0.126 0.171 0.120

66 70 56

Orrett and Karl, 1987

10 50 90

UV-PO

0.14 0.39 0.33

29 35 3

Ridal and Moore, 1990

NE continental shelf slope (George Banks)

Surface 200 400-800

UV

0.17 0.06 0.03

Southern NW shelf, Australia Eel River Shelf, N. California (summer values)

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).

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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., P/, it may be more reasonable to acknowledge the possible role of DOP in microbial metabolism. The TDP model, therefore, assumes that both P/ and DOP are bioavailable during these relatively short ( PP ^- DOP) and four-component (Fi -^ PP1/PP2 -^ DOP) steady-state models to evaluate the observed P pools and fluxes. Their results were consistent with a rapid and coupled production and utilization of DOP in these marine habitats. Dolan et al (1995) examined coupled Fi uptake and passage through a simplified planktonic food web as defined by specific particulate matter size fractions in Villefranche Bay, France. Fi uptake was dominated (>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

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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

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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

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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.

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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.

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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

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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,

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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

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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)

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(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

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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).

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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 <j

-400

MAB Deep Slope Samples \ y

MAB nepheloid layer, >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

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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

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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

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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.

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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.

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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 -4.0

-5.0 Figure 1 Microbial respiration in unfiltered surface water during dark incubation with deck irradiated 0.2-)Ltm-filtered surface water (2 parts unfiltered water to 1 part irradiated filtered water) relative to the unirradiated dark control. Both filtered and unfiltered samples were from a depth of 10 m at a site about 2 km off the mouth of the Delaware Bay, June 2000.

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apparently contained sufficient concentrations of new substrates to maintain high respiration rates even after 12 h of growth. Photodegradation of DOM increased respiration rates by 0.15 /xM CO2 h"^ which is comparable to the abiotic CO2 photoproduction rate measured in the same waters (Kieber etai, 2001). Combining abiotic CO2 photoproduction with photochemically enhanced microbial respiration yielded an increase in the rate of carbon cycling of about 3-5 fiM C day~^ at this coastal station. DOC concentrations of 80-100 /xM C at this site yields DOC turnover times of 16-33 days. These results indicate that photochemical degradation can significantly impact carbon cycling in coastal surface waters. This finding is consistent with results from photobleaching studies in the same sampling region (Vodacek et al, 1997) and photoremineralization studies conducted in DOM-rich (and DOM-augmented) fresh waters and estuarine waters (Miller and Zepp, 1995; Miller and Moran, 1997) and modeling studies in oceanic and estuarine waters based on extrapolation of CO photoproduction rates (Stubbins, 2001). While respiration was clearly enhanced in the above coastal study, microbial production (as measured by both leucine and thymidine incorporation methods) decreased by about 20% in irradiated surface samples relative to dark controls (Mopper et ai, unpublished results). Based on the production results alone, one would conclude that photochemistry negatively impacted microbial carbon utilization. However, when the respiration results are also considered, the overall effect was that photochemistry positively impacted bacterial carbon utilization. Our study also demonstrated that BGE is not constant as often assumed (see review by del Giorgio and Cole, 2000). Instead, the BGE shifted to lower values in the irradiated samples due to enhancement of respiration over production, as observed in some freshwater studies (Vahatalo and Salonen, 1996; Anesio etal, 2000; Faijalla etai, 2001). The photochemical effect on bacterial activity is dominated by its effect on respiration, since respiration generally accounts for the largest fraction of bacterial carbon demand (del Giorgio and Cole, 2000; WiUiams, 2000). Given the overall importance of respiration, it is inappropriate to use bacterial production measurements alone to assess the effect of DOM photodegradation on bacterial activity. These findings also suggest that conclusions from many past uptake studies regarding the net photodestruction of substrates or photochemical formation of biologically recalcitrant humic substances from substrates may need to be reevaluated. A further complication with bacterial production measurements was revealed in glucose addition experiments that addressed carbon limitation. Surprisingly, it was often observed that glucose additions had either no effect or caused a significant decrease in leucine and thymidine incorporation (Fig. 2A), even though glucose was clearly taken up (Fig. 2B). A lack of effect of glucose additions on the growth of planktonic bacteria has been previously observed (Kirchman, 1990; Pomeroy etai, 1995; Carlson et ai, 1999). In the case of decreased incorporation, it is possible that the added glucose stimulated the internal production of leucine and thymidine, thereby diluting the radiolabel, which would result in an apparent decrease in

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Figure 2 (A) Effect of glucose additions on the microbial uptake of radiolabeled thymidine and leucine. Unfiltered samples were from a depth of 10 m at a site about 2 km off the mouth of the Delaware Bay, June 2000. Thymidine and leucine uptake rates were measured according to Smith and Azam (1992) and converted to apparent productivities according to Kirchman (1990) and Kirchman and Rich (1997). (B) Community uptake of uniformly radiolabeled glucose (^"^C) as a function of added glucose (in the same seawater sample as A). The saturation curve describes the uptake kinetics (affinity constant K = 6.49 nM), as given by the equation in the figure.

uptake of radiolabel (P. Falkowski, pers. commun., 2001). A similar effect was found for pyruvate in Antarctic waters (Mopper et al, unpublished results). Since pyruvate is a photochemically produced substrate (Kieber et al, 1989), the question arises as to whether the apparent "inhibition" effects reported in several studies (e.g., decreased microbial production after irradiation; Appendices 2 and 3) may be partly (or totally) due to the photochemical production of substrates that, after being taken up, stimulate the internal production of leucine or thymidine. Clearly, this question needs to be addressed in future studies.

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2. Dissolved Inorganic Carbon Formation and Oxygen Consumption While it has been known for some time that photolysis of DOM yields LMW organic products (see review by Kieber, 2000), it has only recently been shown that the major photoproducts are inorganic species, in particular CO and dissolved inorganic carbon (DIC) (Mopper et ai, 1991; Miller and Zepp, 1995; Bates et al, 1995; Zuo and Jones, 1995; Stubbins, 2001). Photochemical DIC formation may strongly impact carbon cycling in seawater and other natural waters (Miller and Zepp, 1995; Miller and Moran, 1997; Kieber et al, 1999a, 2001; Lindell et al, 2000). In addition, this process may be responsible for a major fraction of abiotic oxygen consumption in irradiated surface waters (Lindell and Rai, 1994). However, photochemical formation pathways of DIC and how they are linked to oxygen consumption are not understood. Evidence supports two main routes for DIC production from DOM: one by oxygen-independent pathways (e.g., direct decarboxylation) and the other by reaction with molecular oxygen and/or reactive oxygen transients (Appendix 4). In support of the latter pathway. Miles and Brezonik (1981) proposed a decarboxylation mechanism that involves the consumption of molecular oxygen, i.e., 1 mol of molecular oxygen consumed per 2 mol of CO2 produced, which is close to the ratio measured in their study. In contrast, Lindell and Rai (1994) and Amon and Benner (1996) found a molar O2 consumption to CO2 evolution ratio close to one for lake waters and river waters, respectively. Despite this close mass balance, it cannot be concluded from these studies that photochemical production of CO2 is dependent on the photoconsumption of molecular O2. Indirect evidence for this dependency comes from Lindell et al (2000), who found no DIC photoproduction from continuously irradiated DOM-rich water samples after reaching 25% or less O2 saturation resulting from photochemical O2 consumption. However, it is not clear whether the drop in DIC production was due mainly to the drop in O2 partial pressure or to the photodestruction (photobleaching) of precursors, as has been demonstrated for CO (Stubbins, 2001). The above studies suggest that molecular oxygen is required for DIC photoproduction. However, this mechanism is not consistent with the finding that DIC photoproduction in sterile filtered coastal seawater either did not change or increased by about 35% after molecular oxygen was removed by exhaustive sparging with oxygen-free helium (Kieber et al, 1999a, 2001). Furthermore, the existing experimental results are not consistent with the amount of O2 required for all purported oxidative processes. For example, about 50 to 100% of all photochemically consumed oxygen is required for the production of H2O2 via dismutation of superoxide (Petasne and Zika, 1987; Blough and Zepp, 1995; Blough and Caron, 1995; Andrews et al, 2000) and for DOM oxidation (Blough and Zepp, 1995). Presumably, this oxygen would not be available for DIC formation. The latter studies suggest that DIC may be photochemically formed from DOM by

Photochemistry and the Cycling of Carbon, Sulfur, Nitrogen and Phosphorus oxygen-free pathways, e.g., by cleavage of carboxyl groups from DOM molecules (decarboxylation). A variety of decarboxylation pathways have been identified using model compounds (Budac and Wan, 1992). Some potentially important photochemical mechanisms include cleavage of a carboxyl group attached to an aromatic ring via an intramolecular or intermolecular charge transfer reaction (Budac and Wan, 1992); ligand to metal charge transfer excitation of metal complexes (or chelates) followed by cleavage of the carboxyl group (Langford et al, 1973); OH radical-induced decarboxylation of amino acids (Steffen et al, 1991) and organic acids (Zafiriou, 1990); photolysis of aromatic rings followed by decarboxylation of ring cleavage products (Chen et al, 1978; Vahatalo et al, 1999); and photolysis of LMW organic acids (Bockman et al, 1996) such as pyruvate, glyoxylate (Kieber, 1988; Klementova and Wagnerova, 1990) and oxalate (Bertilsson and Tranvik, 1998). The latter mechanism may involve photosensitized oxidation (Klementova and Wagnerova, 1990). If Die photoproduction from DOM occurs mainly by direct decarboxylation (as opposed to oxygen-dependent pathways), one might expect DIC photoproduction to be highly correlated to DOM carboxyl content. However, Mopper et al (2000), in a study of DIC photoproduction by humic substances added to fresh water, found that the only strong correlation was with aromaticity and not with total carboxyl carbon. The correlation of DIC photoproduction with aromaticity suggests that carboxyl groups attached to aromatic rings are preferentially cleaved off the ring to form CO2, which is consistent with studies of model compounds (Budac and Wan, 1992). This mechanism is also consistent with IR spectroscopic results for fulvic acids that showed that the carboxyl signal associated with aromatic groups is preferentially lost during UV irradiation of aqueous solutions, while the aliphatic carboxyl signal increased and the fulvic acids became more oxidized and less aromatic (Chen et al, 1978). Similar results were obtained by solid-phase ^^C and two-dimensional NMR analyses of photolyzed aqueous fulvic and humic substance solutions (Kulovaara et al, 1996; Schmitt-Koplin et al, 1998). Thus, the above studies further support the idea that photochemical DIC production is largely independent of photochemical molecular oxygen consumption. It is clear from these conflicting reports that the major pathways involved in DIC photoproduction, and their relationship to DOM photooxidation and oxygen consumption in seawater and other natural waters, need further study. B, SULFUR Organic sulfur compounds are present in seawater at low concentrations, typically in the 10~^^-10~^ M range, due to close coupling between production and removal processes that include uptake and release by biota, thermal reactions, and photochemical reactions. Photochemical transformations are a key component

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of the sulfur cycle in the upper ocean, and an important factor controlling the volatility and efflux of sulfur from the oceans into the atmosphere. In the atmosphere, reduced sulfur species are oxidized to photochemically stable species including sulfuric acid and methane sulfonic acid, a process that results in the formation of cloud condensation nuclei (CCN), which in turn affect the Earth's radiative balance (Charlson ^r a/., 1987). The principal trace sulfur compounds in seawater are dimethyl sulfide (DMS), dimethylsulfoniopropionate (DMSP), carbonyl sulfide (OCS) and dimethyl sulfoxide (DMSO). Other minor species include hydrogen sulfide (H2S), dimethyl disulfide (DMDS), carbon disulfide (CS2), methane thiol (MeS), cysteine, glutathione, phytochelatins and methionine. Many of these compounds are either formed or degraded in seawater through photochemical transformations involving DOM. However, except for disulfides, none of these sulfur compounds absorb solar radiation in seawater and, therefore, they will either be photochemically stable or will photochemically degrade through secondary photochemical pathways involving photosensitizers and/or reactive transients. Sulfur compounds undergo a variety of secondary photochemical transformations in aqueous media, especially alkyl sulfides, disulfides, and thiols. Known reactants for sulfur include the hydroxyl radical, hydrogen peroxide, and singlet oxygen. However, none of these species is quantitatively important in sulfur phototransformations in seawater, as discussed below. 1. Dimethyl Sulfide Dimethyl sulfide is the predominant volatile sulfur species found in surface seawater (Kettle et al, 1999a). Its atmospheric loss and subsequent photooxidation plays an important role in the formation of CCN in the troposphere (Charlson et al, 1987; Andreae etal, 1995). The primary source of DMS is through enzymatic lysis of DMSP, although it is not clear why algae ly se DMSP (Stefels, 2000; Sunda et al, 2001). While DMSP undergoes enzymatic lysis, it does not photolyze in seawater at ambient concentrations (ca. 10~^ M) (Kieber and Kiene, unpublished results). Atmospheric ventilation was originally proposed to be the primary removal pathway for DMS in seawater, but now it is clear that microbial uptake and photochemical degradation often control the removal of this compound in the photic zone (Kieber et al, 1996; Dacey et al, 1998; Simo and Pedros-Alio, 1999). Dimethyl sulfide does not absorb solar radiation in seawater, and therefore its photochemical degradation must involve other photochemically active components in seawater such as chromophoric DOM. In support of this supposition, the photochemical degradation of DMS is highly correlated to the concentration of DOM (Jiao and Kieber, 1996; Brugger et al, 1998), but the role of DOM in the photochemical breakdown of DMS is not known. The photochemical loss of DMS in seawater is pseudo-first-order with respect to DMS concentration (Brimblecombe and Shooter, 1986; Kieber et al, 1996;

Photochemistry and the Cycling of Carbon, Sulfur, Nitrogen and Phosphorus Brugger et al, 1998). However, at high concentrations of added DMS, observed reaction kinetics approach zero order with respect to DMS. This saturation-type behavior suggests that the photochemical loss of DMS occurs through a binding (or catalytic) mechanism, presumably involving components of DOM (Jiao and Kieber, 1996) and perhaps reactive species that are generated by DOM. The wavelength dependence for the photochemical loss of DMS determined in the equatorial Pacific showed a distinct maximum in the blue at approximately 450 nm (Kieber^r al, 1996). However, because of the extremely low absorbance of this seawater, it was not possible to calculate apparent quantum yields above 400 nm. The maximum at 450 nm is presumably due to a specific, photoactive component in seawater, probably of biological origin, which is not always present in the seawater. Indeed, further wavelength dependence studies in other oceanic waters did not always show the presence of this peak (Kieber et al, unpublished results). Photochemically generated singlet oxygen (^Oi) can be an important oxidant at micromolar levels of DMS (Brimblecombe and Shooter, 1986). However, it is a relatively minor oxidant at ambient DMS and ^02 concentrations, accounting for about 14% of total DMS photochemical loss, as determined from direct DMSO measurements and sodium azide ^02 quencher studies (Kieber et al, 1996). This finding is consistent with calculations that yield a DMS removal rate (or DMSO production rate) of 2.1 x 1 0 - i i M h - \ given 1 x 10-^MDMS,1 x 1 0 - I 3 M 1 O 2 , and a bimolecular rate constant of 5.8 x 10^ M~^ s~^ (Wilkinson et al, 1995). As with singlet oxygen, hydroxyl radical concentrations (ca. 10~^^M) are also too low to effectively remove DMS from seawater, even though the bimolecular rate constant for this reaction is near the diffusion-controlled limit (k= 1.9 x 10^^ M"^ s~^; Buxton et al, 1988). Dimethyl sulfide also reacts with hydrogen peroxide in seawater (ca. k = 0.14 M~^ s~^; Shooter and Brimblecombe, 1989), but at ambient concentrations of DMS and H2O2 (ca. 10"^ and 10~^ M, respectively), reaction rates are too slow (ca. 0.1 x 10~^^ M h~^) to be a significant removal mechanism for DMS in the photic zone. Appreciable rates are expected only when high concentrations of DMS and/or H2O2 are encountered such as may be observed inside an algal cell (Sunda et al, 2001), or possibly during a rain event (Cooper et al, 1987), or under bloom conditions when relatively high H2O2 and DMS concentrations may occur in the water column. Thus, although it is evident that DOM is involved in the photosensitized loss of DMS in seawater, none of the known reactions can account for the observed DMS loss when considered alone or together. Furthermore, the predominant products have not been identified, as DMSO represents less than 15% of the total DMS lost. 2. Dimethyl Sulfoxide Dimethyl sulfoxide is released into seawater through biological processes (Lee et al., 1999), and is produced in seawater through the reaction of DMS with singlet

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oxygen (Brimblecombe and Shooter, 1986; Kieber etal, 1996). Dimethyl sulfide may also react with other reactive oxygen species such as the carbonate radical, which is known to react with organic sulfides (Huang and Mabury, 2000), to form DMSO. Dimethyl sulfoxide cannot undergo direct photolysis in seawater, since it does not absorb actinic radiation (ca. >290 nm). Therefore, if DMSO photochemical loss is observed, it must occur through secondary photochemical pathways. Reaction of DMSO with the hydroxyl radical is too slow to be quantitatively important. Given a bimolecular rate constant of 6.6 x 10^ M~^ s~^ (Buxton et al, 1988) and ambient seawater concentrations of DMSO and the OH radical of 25 X 10"^ M and 1 X 10"^^ M (Mopper and Zhou, 1990; Vaughan and Blough, 1998), respectively, the rate of DMSO loss in surface seawater is only 6 X 10"^^ M h~K Significant rates are expected only at higher concentrations of reactants, which may be encountered, for example, in a marine algal cell containing high DMSR Inside the cell, DMSO may play a role in alleviating cellular OH radical stress (Lee and de Mora, 1999; Sunda et ai, 2001). Based on rate calculations with known reactants and observations by Brimblecombe and Shooter (1986), it is unlikely that DMSO is photochemically degraded in seawater as suggested by Lee et al (1999). Photochemical studies with ambient levels of radiolabeled DMSO are needed to address this question. While there is some uncertainty regarding the photochemical loss of DMSO, it is likely that this nonvolatile compound will be removed from seawater primarily through its microbial uptake, although experimental evidence to support this removal pathway is limited (Kiene and Gerard, 1995; Lee ^r a/.,'1999). 3. Carbonyl Sulfide Carbonyl sulfide is the most stable, naturally occurring sulfur species that is ventilated from the oceans into the atmosphere. It has a lifetime in the troposphere of more than 1 year (Khalil and Rasmussen, 1984) and diffuses into the lower stratosphere, where it is photooxidized to sulfuric acid, giving rise to the Junge aerosol layer (Crutzen, 1976). In seawater, concentrations of OCS are low compared to those of DMS, DMSO and DMSP, ranging from approximately 0.071.2 X 10"^ M in coastal waters to about 0.03 x 10~^ M in the open ocean. Low seawater concentrations of OCS reflect its low production rate in seawater and rapid turnover in the water column. In particular, although OCS is relatively longlived in the troposphere, it is very reactive in seawater due to its rapid hydrolysis to carbon dioxide and hydrogen sulfide (EUiot et al, 1987,1989; Flock and Andreae, 1996). The half-Hfe of OCS in seawater is approximately 2 days ( k = 3.8 x 10"^ s~^) at 298 K. Photochemical loss of OCS is unlikely to be a major sink as OCS does not undergo primary photolysis in seawater and its reaction with known photochemically formed oxidants is very slow, especially when compared to other organic sulfur compounds.

Photochemistry and the Cycling of Carbon, Sulfur, Nitrogen and Phosphorus Carbonyl sulfide is produced in the photic zone by both thermal and photochemical pathways. Nonphotochemical (thermal) production of OCS involves DOM (Flock et al, 1997), but the mechanism is poorly described. Nonetheless, thermal formation of OCS is an important source for this compound in seawater, comparable to photochemical production rates (Flock and Andreae, 1996; Ulshofer et al, 1996). Nonphotochemical production is presumably the main source of OCS in deep seawater. Photochemical production is an important source of OCS in sunlit surface waters (Ferek and Andreae, 1984), with rates in the 10"^^ M h~^ range (Flock and Andreae, 1996). Photochemical production of OCS involves DOM both as a photosensitizer and as a source of sulfur (Andreae and Ferek, 1992). Mechanistic studies have shown that oxygen-dependent oxidants such as singlet oxygen and the OH radical are not involved in OCS formation (Zepp and Andreae, 1994). A number of model sulfur compounds can generate OCS photochemically, including cysteine, 3-mercaptopropionic acid, glutathione, thiols, sulfides, and DOM itself, while DMSP, DMSO, and dimethylsulfone produce very little OCS (Zepp and Andreae, 1994; Flock et al, 1997). Production rates increase when DOM is added as a photosensitizer (Flock et al, 1997; Uher and Andreae, 1997). Based on these and other published studies, the mechanism for OCS photoproduction is proposed to involve thiyl or sulfhydryl radicals (Flock etal, 1991 \ Pos etal, 1998). However, the source(s) of these radicals is not known. Gun et al. (2000) studied polysulfide formation in Lake Kinneret, and suggested that OCS formation results from the reaction of polysulfide radicals and carbon monoxide. This mechanism may explain the strong correlation between OCS photoproduction and CO concentrations observed in marine waters (Flock et al, 1991 \ Pos et al, 1998). Perhaps the source of thiyl radicals is the photolysis of metal-sulfide complexes through a ligand to metal charge transfer reaction. These complexes are known to occur in seawater (Luther and Tsamakis, 1989), but their stabiHties, especially with respect to photolysis, are not known. Apparent quantum yields for OCS photoproduction decrease exponentially with increasing wavelength in the UV (Zepp and Andreae, 1994; Weiss et al, 1995), yielding a broad solar response curve between 290 and 400 nm with a maximum response centered at approximately 340 nm (Weiss et al, 1995). This trend is similar to that observed for many photochemically generated species in seawater (see Blough, 1997, for review). However, apparent quantum yields for OCS production (ca. 10~^) are generally much lower than observed for other species (ca. lO""*10~^), such as CO, CO2, and hydrogen peroxide, indicating that specific (and minor) components of DOM are precursors to OCS. Apparent quantum yields for OCS photoproduction were generally much higher in the Zepp and Andreae (1994) study in coastal waters (e.g., 7.0 x 10~^ at 330 nm) than those obtained in open oceanic waters by Weiss et al (1995)(e.g., 1.6 x 10~^ at 336 nm). Other than analytical artifacts, the only way to generate these large differences is to invoke

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large differences in the concentration or type of sulfur precursors present in those samples. If OCS apparent quantum yields vary greatly in seawater, then it may be difficult to model water colunm production rates on a global scale, especially when using remotely sensed data for DOM fluorescence or absorbance (Neale and Kieber, 2000). 4. Minor Suljfur Species Carbon disulfide, methane thiol and hydrogen sulfide, present in seawater at 10~^^ M levels, are all involved in photochemical transformations (vide infra). Other reduced sulfur compounds that may also be involved in photochemical transformations in oxic seawater, but which have not been studied, include cysteine, cystine, glutathione, and polysulfides. Carbon disulfide, which is an atmospherically important trace gas, is photochemically produced in surface waters. Apparent quantum yields for CS2 formation decrease exponentially with increasing wavelength from about 14 x 10~^ at 308 nm to 0.2 x 10~^ at 400 nm (Xie et al, 1998). Apparent quantum yields for CS2 formation are orders-of-magnitude lower than those of hydrogen peroxide or Die and they are approximately 25% of corresponding apparent quantum yields for OCS formation (Weiss et ai, 1995). Both cysteine and cystine are efficient precursors of CS2, and the OH radical is likely an important reactant (Xie et al, 1998). Based on field and laboratory evidence, Xie et al (1998) concluded that the mechanisms for CS2 and OCS formation are similar, and that photoproduction was the primary source of CS2 in seawater. Methane thiol and hydrogen sulfide are both removed from seawater through photochemical transformations. Average methane thiol removal rates in the Northeast Atlantic and Aegean Sea were low, 6,7 and 36 x 10"^^ M h ~ \ respectively (Flock and Andreae, 1996). Although the mechanism is not known, the photochemical loss of methane thiol involves binding to DOM possibly through complexation or formation of a thioketal (Kiene et al, unpublished results). The photochemical degradation of hydrogen sulfide followed first-order kinetics in Biscayne Bay, Rorida seawater, with loss rates of 5.1 and 20.2 x 10~^^ M h"^ Photochemical loss of sulfide was independent of dissolved oxygen and involved DOM and possibly a thiyl radical intermediate (Pos et al, 1997). These investigators observed that the rate of sulfide loss did not decrease exponentially with increasing wavelength as observed for most other species in seawater. Rather HS~ loss showed a maximum in the blue region of the solar spectrum (Pos et al, 1997), which was also observed for DMS (Kieber et al, 1996). It is clear from the above studies that photochemical transformations affect many reduced sulfur compounds in seawater, but there is still much that is not known. In particular, the role of iron and other metals in complexing sulfur compounds, especially hydrogen sulfide and thiols, in seawater is poorly understood. Sulfur

Photochemistry and the Cycling of Carbon, Sulfur, Nitrogen and Phosphorus compounds complex very strongly to type B metals, such as iron and copper, to form very strong complexes. Thus, metallo-sulfur complexes are likely to be important in seawater, both in stabilizing sulfur to thermal oxidation and destabilizing sulfur to photochemical oxidation through ligand to metal charge transfer reactions. Other less-known reactants may also be important in sulfur photodegradation in seawater including the carbonate radical, superoxide anion and organic radicals. Systematic studies are needed to quantify and delineate these sulfur phototransformations.

C. NITROGEN AND PHOSPHORUS In comparison to carbon or sulfur, relatively little is known about photochemical transformations of nitrogen and phosphorus in natural waters, and the available information indicates that these transformations are complex. This complexity is evident from the contradictory results that have been reported to date. Bushaw et at. (1996) irradiated whole water or isolated fulvic acid extracts from a boreal pond and a series of high DOM rivers in Georgia (including the Satilla River), with natural sunlight (or artificial light), and measured high rates of NH4" production, ranging from 40 to 370 x 10~^ M h" ^ Photochemical NHJ production was also observed in irradiated Satilla River water employing specific detection of ammonium by HPLC (Kieber, 2000). Gardner et al (1998) conducted a i^NH+ isotope dilution study infilter-sterilizedlake water and found photoproduction rates for ammonium (ca. 200 x 10~^ M h~^) comparable to those obtained by Bushaw et al. (1996). Additionally, Gardner et al. (1998) were also able to detect variable but consistent photochemical loss of the anmionium (2-130 x 10"^ M h"^) in the same water, presumably due to either incorporation into DOM or loss through a secondary photooxidative pathway. The latter finding is consistent with the results of a laboratory study by Kieber et al. (1997), which demonstrated that ammonium is incorporated into humic extracts during photolysis of filtered seawater. The mechanism for ammonium photoproduction (or loss) was not investigated in these studies, but it may involve either the release of bound ammonium or the photochemical breakdown of DOM to form ammonium (e.g., deamination of peptides or amino acids). Kinetic results support the second hypothesis (Zepp, unpublished results). Although a number of studies have observed net photochemical production of ammonium formation in natural waters, there are many reports that show either no photoproduction or a photoinduced loss of NH^ (e.g., Koopmans and Bronk, 2002). Vahatalo and Salonen (1996), J0rgensen et al. (1998), Bertilsson et al. (1999), and Wiegner and Seitzinger (2001) all observed no change in ammonium concentrations when filter-sterilized fresh water was exposed to sunlight, while J0rgensen etal. (1999) observed a 35-82% loss (or no change) of NHj in irradiated

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Baltic Sea water samples. The reasons for these contrasting results are not known, but may be due to inherent differences in the DOM composition among the different water samples and to differences in the relative importance of photoproduction versus photodestruction pathways for anmionium, as suggested by the work of Gardener et al (1998). Alternatively, the contrasting results may reflect differences and limitations in the analytical methodologies used to quantify NH^, or differences in light exposure history of the DOM (Koopmans and Bronk, 2002). As with ammonium, photochemical results for primary amines and amino acids are not consistent (Koopmans and Bronk, 2002). Using a batch fluorometric technique, Bushaw et al (1996) and Bushaw-Newton and Moran (1999) reported photoproduction of primary amines in DOM isolated from the Skidaway and Satilla River samples. By contrast, Kieber and Miller (unpublished results) found no chromatographic (HPLC) evidence for primary amine or amino acid production in a Satilla River fulvic acid isolate above the detection hmit (ca. 10"^^ M), even after 11 h of solar irradiation. J0rgensen et al (1998) also employed HPLC, but in their study, dissolved free amino acids (DFAA) and carbohydrates increased when water from Lake Skarshult was exposed to sunlight. However, the increases in DFAA were small (13-23%), and no information was given regarding changes in specific amino acids. Increases in DFAA concentrations are likely due to their photochemical release from organically bound DOM. In a follow-up study, J0rgensen et al (1999) observed that in some irradiated samples DFAA concentrations increased, while for other samples there was a net decrease (or no change) in DFAA concentrations. Again, no specific compositional changes in the DFAA pool were reported. In a study assessing the effect of sunlight on protein lability, irradiation of seawater with added protein apparently made the protein less available to microorganisms, possibly due to production of refractory DOM with a concomitant loss of labile proteinaceous nitrogen (Keil and Kirchman, 1994). Thus, photoreactions involving DOM may be both a source and a sink of biologically available nitrogen. In related studies, Amador et al (1989,1991) demonstrated that the photolysis of humic-bound organics such as amino acids and aromatic compounds greatly increases the rate and extent that microorganisms degraded these compounds relative to dark controls. They found that only 20% of the ^"^C-labeled glycine in their humic sample was utilized over a 60-day period in the dark. Most of the glycine (ca. 80%) was bound to intermediate and high-molecular-weight humics (>5000 Da) that could not be utilized by the microbes. Following the exposure of humic acid solutions to sunlight, the percentage of glycine that was mineralized by the heterotrophic bacteria in the dark increased to 40-60% of the total. This increase paralleled the increase in the percentage of glycine associated with the LMW humic acid fraction, and is consistent with the finding that only the photochemically formed LMW fraction is mineralized by the bacteria. However, not all humic-bound organics are biologically labile after exposure to sunlight. They

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attributed this incomplete utilization to the formation of photochemically and biologically resistant iV-heterocycles in the humic acids (Amador et al, 1989, 1991). The lack of agreement in DFAA and NHj results has also been seen for other nitrogen compounds. Photoproduction of urea and nitrite have been observed in natural waters (J0rgensen^r a/., 1998, \999\YAthQv et al, 1999b; Jankowski^ra/., 1999; Koopmans and Bronk, 2002). These compounds have also been shown to remain constant or decrease when filter-sterilized samples are exposed to sunHght (Vahatalo and Salonen, 1996; J0rgensen et al, 1999; Kieber et al, 1999b; Jankowski et al, 1999). For nitrite, these differences are expected because nitrite undergoes primary photolysis in seawater (Zafiriou and True, 1979; Zafiriou and Bonneau, 1987) that competes with its photoproduction. In coastal seawater, the rate of direct photolysis of nitrite was approximately 2.3 x 10~^ M h~^ compared to an average photochemical production rate of 4 x 10~^ M h~^ (Kieber et al, 1999b). Photochemical formation of phosphate was first reported by Francko and Heath (1982), who observed that sunhght exposure of DOM-rich, acidic lake water (pH 5.2-5.8) caused the release of phosphate complexed to high-molecular-weight DOM. The rate of release of bound phosphate was shown to be highly correlated to reduction of Fe (III) to Fe (II). Phosphate photoproduction was also observed in a humic-rich lake (Vahatalo and Salonen, 1996). However, the mechanisms underlying phosphate photoproduction are poorly understood (Francko, 1990), and there appears to be a seasonality in the extent of phosphate release (Cotner and Heath, 1990). Photochemically induced increases in dissolved phosphorus levels have not been observed in other systems, including a culture of the marine diatom A. anophagefferens (Gobler et al, 1997) and a freshwater lake sample (J0rgensen^rfl/., 1998). It is somewhat surprising that DOM shows such a diverse photochemical response in natural waters with respect to nitrogen and phosphorus (i.e., no production, production, and loss). It is difficult to reconcile these differences because the photochemical mechanisms underlying these transformations are not understood, even at the most basic level. For example, what is the effect of dissolved oxygen, pH, and temperature on production rates? Are production rates linearly dependent on the photon exposure? What is the observed reaction order? What are the wavelength-dependent apparent quantum yields? Conflicting results may be due to fundamental differences in the DOM among these waters or to variations in the balance between rates of photoproduction and photodestruction. They may also reflect differences in experimental design and analytical artifacts that result from a lack of understanding of the factors that control nitrogen and phosphorus phototransformations in natural waters. Given the wide range of photochemical results that have been obtained for nitrogen and phosphorus, it is difficult to evaluate the overall importance of photochemical transformations on their biogeochemical

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cycling. This is especially true in oligotrophic seawater where no studies have been performed, almost certainly due to analytical limitations.

IV. UNRESOLVED QUESTIONS AND FUTURE RESEARCH Photochemical reactions undoubtedly have an important impact on a variety of marine processes, and potentially play a significant role in the global biogeochemical cycles of carbon, nitrogen, sulfur, and perhaps phosphorus. However, even though there has been an exponential increase in the number of marine photochemical studies over the past decade, many important questions remain unanswered. We conclude this chapter by indicating future research areas, which include developing a mechanistic understanding of photochemical processes, identifying and quantifying the major photochemical fluxes, and extrapolating/ modeling these fluxes to global scales.

A. MECHANISTIC STUDIES

Despite the importance of photochemistry in the surface ocean, relatively few comprehensive mechanistic studies have been carried out, and those that have been conducted have concentrated on identifying reactive transients and on measuring their steady-state concentrations, rate laws, and production rates in seawater (Blough andZepp, 1995; Blough, 1997; Faust, 1999). Details of most photochemical reaction pathways have not been determined and virtually nothing is known about the major reactive sites (or chromophores) within marine DOM that are responsible for the production of reactive transients in the sea. Progress in this area has been inhibited by the lack of sensitive and selective probes suitable for seawater, as well as a lack of studies that track element mass balances resulting from photochemical transformations. For example, can the photochemical loss of DOC in seawater be accounted for mainly by the production of DIC (and CO)? In order to develop a predictive understanding of the rates and controls of important photochemically produced compounds, it will be necessary to identify the DOM precursors and photosensitizers that affect specific photochemical reactions (i.e., which chromophores are involved?). It will also be important to determine the factors (in chemical, biological and physical domains) that control the photoproduction/ destruction of key species (e.g., DIC, CO, DMS). For example, what is the role of O2 in DOM photodegradation and DIC production; i.e., is molecular oxygen needed as a reactant or does it suppress DIC photoproduction by, for example, quenching reactive triplets, or both? What is the fate of the resultant photooxidized DOM (Blough and Zepp, 1995)? What are the mechanisms and major products formed during the photolysis of DMS in seawater?

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B. PHOTOPRODUCTION AND AIR-SEA EXCHANGE OF IMPORTANT ATMOSPHERIC TRACE GASES Climate models are greatly affected by the degree to which air-sea gas exchanges of CO, C O S , D M S and other cHmatically important gases are influenced by photochemical reactions at the sea surface (Doney et al, 1995; Najar et aL, 1995; Gnanadesikan, 1996; Blough, 1997; Erikson et al, 2000). Future studies are needed to quantify the photoproduction of CO2 and photoproduction/loss of volatile sulfur species in surface seawater and to determine the effect of these processes on their air-sea exchange, as has been done for other reactive species in the sea surface microlayer (Thompson and Zafiriou, 1983). Furthermore, the impact of changes in global UV radiation on air-sea fluxes of important trace gases will need to be assessed. Another major unknown is the role of photochemical processes in the formation and/or destruction of the surface microlayer and its impact upon air-sea gas exchange (Blough, 1997).

C. DOM

PHOTOCHEMISTRY AND MARINE FOOD WEB DYNAMICS

The photochemical processes by which biorefractory DOM is made bioavailable, and by which bioavailable DOM is made unavailable, need to be determined. Identifying these processes would help in developing a predictive understanding of changes in DOM lability due to coupled photochemical-biological transformations. In this context, quantifying the effect of photochemistry on bacterial growth efficiency appears to be essential. In a related area, the role of extracellular release of DOM as a photoprotective mechanism in autotrophs and heterotrophs needs to be examined; i.e., what is the effect of externally (and internally) produced reactive phototransient species (e.g., free radicals) on the growth of autotrophs and heterotrophs?

D. DOM PHOTOCHEMICAL REACTIONS INVOLVING TRACE ELEMENTS Photochemistry can affect trace element speciation (Blough and Zepp, 1995; Faust, 1999). For example, photochemical reactions can alter the concentration and reactivity of organic ligands involved in the complexation/solubilization of trace metals (Moffett, 1995). Photochemical reactions can also result in changes in the oxidation state of trace elements, which in turn can affect their bioavailability. Redox-sensitive trace elements include iron (Voelker et al, 1997; Barbeau and Moffett, 2000), copper (Zafiriou et al, 1998; Voelker et al, 2000), manganese (Sunda and Huntsman, 1994), chromium (Kaczynski and IGeber, 1993),

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mercury (Costa and Liss, 2000; Lalonde et al, 2001) and iodine (Wong and Cheng, 2001). The mechanisms by which these photoredox reactions occur, and the nature of the DOM ligands involved, are not known, and are clearly areas for future research. While DOM photochemistry significantly impacts the chemistry of important trace metals in seawater, the reverse, i.e., the effect of trace metals on the photochemical reactivity and degradation of DOM, is not understood. What is the fate of the oxidized DOM product? Is its complexing capacity increased or decreased relative to the reduced DOM? Does oxidized DOM undergo further reactions leading eventually to decarboxylation, which in turn would contribute to photochemical DIC production?

E. EXTRAPOLATING PHOTOCHEMICAL RATES TO GLOBAL SCALES In order to evaluate the impact of DOM photochemistry on climatic processes and biogeochemical cycling of elements, it will be important to use remotely sensed estimates of chromophoric DOM (CDOM) and available irradiance data to extrapolate photochemical fluxes to global scales (see Blough and Del Vecchio, Chapter 10). However, modeling will require laboratory studies to develop robust parameterizations for apparent quantum yields for important photochemical reactions, including CO2 photoproduction, H2O2 photoproduction, DMS photodestruction, and CDOM photobleaching (Sikorski and Zika, 1993; Kettle et al, 1999b; Vahatalo et al, 2000). As a complement to this laboratory effort, in situ drifter studies need to be conducted in the field to verify global photochemical models (Zafiriou, pers. conmiun.). The advantage of the latter technique is that it does not have some of the inherent uncertainties associated with models based on apparent quantum yield and absorbance data (Neale and Kieber, 2000). The main drawback to this latter approach is that it is time-consuming and expensive. In addition, seasonal and geographic variability in photochemical reaction rates need to be measured. With this information, it should be possible to estimate the importance of marine photochemical reactions on tropospheric budgets of important trace gases (DMS, CO, COS, etc.). These global estimates can then be compared with estimates of other biogeochemically relevant species and for the development of a unified framework for quantifying global fluxes of carbon and sulfur.

ACKNOWLEDGMENTS We thank the National Science Foundation Office of Polar Programs (OPP-9527255 (K. M.) and OPP-9610173 (D. J. K.)) and the Chemical Oceanography Program (OCE-9711206 (K. M.) and OCE-9711174 (D. J. K.)) for financial support of this work. We thank two anonymous reviewers for constructive criticism of the manuscript, and we also acknowledge Dave Seigel and Norm Nelson

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for valuable comments on the Unresolved Questions and Future Research section. We thank Brendon Hofsetz for collection of the respiration data for Fig. 1 and for assistance with preparation of the figures. We thank Yi Zhu for collecting the uptake data for Fig. 2, and Aron Stubbins for conmients on the page proofs.

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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

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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,

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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

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28

30 32 salinity ( V J

34

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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

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11

0.027 \ 0.024 '

C/5

o

0.021 \

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

0.018 0.015 J

1

•

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MAB mixed layer MAB below mixed layer

o

o

460-

• • •

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V

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•

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420 • 400 J 380 ' 360 1

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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

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r in

r

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•

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-1

60

80

100 120 DOC (/xM C)

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1

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140

.

o (d

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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" - =^

200 m) unless there is a strong downwelling. More realistic can be the transfer through the association with particles. The observation that flocculants can form from dissolved (or, more probably, colloidal) organic matter is not recent but if applied to surface waters enriched with polysaccharides it becomes a possible transfer of important quantities of carbon. The example of the mucilage formed in the Adriatic Sea from DOM, though being an extreme case, describes the possible mechanism. The richness of marine snow in carbohydrates has been mentioned and the bacterial activity associated with these aggregates has been well described (Azam et al, 1993). Recently, a study of the composition of surface and deep water particles has shown that polysaccharides associated with deep-sea particles have a composition close to polysaccharides in surface-water DOM, but different from dissolved sugars found in deep-sea water (Aluwihare and Repeta, 2000). This also suggests that there is an aggregation or adsorption onto particles of recently released carbohydrates in surface water, and that these particles sink rapidly, transferring their signal to deep waters. Another way of DOC transfer has been shown recently (Alldredge, 2000) in the form of interstitial water within marine aggregates. The author found as much as 10-140 mg C L~^ of aggregate, representing one to two orders of magnitude more than in surrounding water. Nevertheless, this "interstitial" carbon represented less than 2.5% of DOC in the water column, which would not change very much the total vertical flux. There is also a high probability that a large fraction of this interstitial DOC comes from the hydrolysis of particulate matter by bacteria in the aggregate and not from DOM from surface water. As far as deep water is concerned, the possibility of surface DOC being aggregated or adsorbed on particles to participate in a rapid vertical flux is probably the most realistic and quantitatively important mechanism of DOC transfer. Considering that seasonal accumulation of DOC in surface waters can increase the DOC concentration by 20 to 200 /xM, with probably a mean increase of 40-50 /xM (about 50 ^lg L"^), (Kepkay et aU 1993; Zweifel et al, 1995; Thingstad et al, 1997; Zweifel, 1999; S0ndergaard etal, 2000; Cauwet etal, 2002) the total DOC increase for the world ocean would be enormous. Considering the different values calculated for yearly DOC accumulation in different oceans (Copin-Montegut and Avril, 1993; Carlson et al, 1994; Borsheim and Myklestad, 1997; see Hansell,

600

Gustave Cauwet

Chapter 15, for global estimates), 1.2 mol C nT'^ would be an acceptable average value. If only 10% of this excess is transferred to deep water it would represent an annual flux of 1.4 g C m~^ year~^ This flux is not very important compared to total primary production but not so different compared to the flux of detrital carbon and will enrich deep waters with labile or semilabile organic carbon. If it is not recycled, it will constitute an additional sink for carbon; if it is recycled it will increase the consumption of oxygen and will influence the cycle of nutrients. The mineralization of DOM generally recycles nutrients (N and P) but if this DOM is mainly composed with carbon-rich molecules (polysaccharides) it will consume more than recycle nutrients, especially phosphorus. This flux is more important in the coastal ocean than in the open ocean due to shallow waters and it will constitute a complementary input of food for some benthic organisms. The lateral export of DOM from the coastal zone to the open ocean should also be considered. Probably because blooms are larger in the coastal zone, the DOC accumulation is more important here than in the open sea. A fraction of this excess DOC can be transported to the slope and the open sea if the dynamics are favorable for such a rapid transfer. The evidence of a filament of carbonrich shelf water abutting the Gulf Stream (Bates and Hansell, 1999) led the authors to calculate that this mechanism could export to the Atlantic as much as 2.7 X 10^^ g C year"^ They also suggest that this DOC originates mainly from riverine sources. It seems far from negligible compared to the estimate of global DOC inputs by world rivers (250 x 10^^ g C year"^). It is difficult to imagine that the Chesapeake Bay exports more than 14% of total DOC exported by the Amazon River (19 x 10^^ g C year"^). It is also surprising that the authors attribute this DOM to the river inputs while the C:N ratio of DOM ranges from 10 to 14, a low value for terrestrial organic matter. Nevertheless, it demonstrates that rapid DOC export from the coastal zone can be an important transfer mechanism.

V. CONCLUSIONS From river inputs to export of DOM, this chapter covers a broad topic and demonstrates in itself the complexity, but also the importance, of the coastal zone. Compared to the open ocean, the nearshore ocean is more complex because of variable inputs from continents, remobilization processes and a succession of several recycling steps. In the coastal ocean, mixing of organic matter of various origins and of different diagenetic histories is the norm. In terms of riverine inputs, most global estimates are more than 10 years old. There is also very little work on the nature of riverine organic matter and its variability. Usually, the viewpoint in the Hterature is that terrestrial organic matter originates from higher plants, forgetting that there is an intense aquatic life in

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freshwater bodies. Some isotopic studies on organic matter exported by estuaries or salt marshes have shown that the terrestrial signature is not always dominant and the aquatic contribution can modify the 8^^C of DOM (Fry et al, 1992). It could explain why some studies on isotopic composition in coastal zone and estuaries give sometimes results difficult to interpret. What is probably true for POC is not certainly true for DOC. The paradox several times mentioned that terrestrial organic matter brought by the rivers is highly refractory to degradation but is rapidly disappearing when progressing toward open sea is not clearly explained. The isotopic measurements made by Kelley et al. (1998) are quite demonstrative of the complexity of coastal systems. These authors found 5^^C values of —20/—25 in particulate organic matter from Mississippi plume to the Gulf of Mexico, -19.6/-24.7 in DOM of the Gulf of Mexico, and - 3 3 in bacteria of the coastal zone, suggesting they are mainly living on terrestrial dissolved organic matter. More recently, Raymond and Bauer (2001) have shown the variable age (from modem to 1300 years) of DOM transported by several rivers of the eastern coast of the United States, compared to the generally old POC (700 to 4700 years). They also demonstrated that the most labile fraction is also the youngest one. The main consequence is the strong influence of the degradation processes occurring in the river, the estuary or the coastal zone on the apparent age of the exported DOM, the youngest one disappearing rapidly while the oldest one can accumulate in the water column or in sediment. Concerning the biogeochemical processes in mixing zones of estuaries, we slowly discover, while progressing in our knowledge that things are much more complex than we thought before and that supposed rich (in terms of nutrients) areas function from time to time like oligotrophic environments. The complex relationship existing between different nutrients and between phytoplankton and bacterioplankton makes interpretation of field results sometimes very difficult. Coupling of fieldwork and experimental approaches will be more and more necessary if we want to understand processes and quantify them. In the future, we must organize new multidisciplinary programs. Studies focusing on a few parameters may be necessary to enrich the data bank but they are too limited for understanding the complex biogeochemical mechanisms occurring in the coastal zone. Until now, for the coastal zone, the water dynamics aspect was somewhat neglected while it could sometimes bring key information. This is particularly true if we consider that one of the most—or maybe the most—important parameter controlling the biogeochemistry of estuarine and coastal zones could be the residence time. The time scale seems to be our bigger problem for explaining the mechanisms. Finally, joining some other authors (Bauer and Druffel, 1998), we must now acknowledge that the coastal zone must be considered seriously as an important contributor to carbon export toward the open and deep sea. To know if it can contribute to a significant carbon sink from the atmosphere to the deep ocean, or even more to the sediment, is another story.

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ACKNOWLEDGMENTS I thank three anonymous reviewers for their useful comments and criticisms, with a special mention for Dr. Jon Sharp, who did extensive editing of my English while also proposing relevant improvements to the manuscript.

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Pettine, M., Patrolecco, L., Camusso, M., and Crescenzio, S. (1998). Transport of carbon and nitrogen to the Northern Adriatic Sea by the Po River. Estuarine Coastal Shelf Sci. 46,127-142. Prahl, F. G., and Coble, P. G. (1994). Input and behavior of dissolved organic carbon in the Columbia River estuary. In "Changes in Fluxes in Estuaries: Implications from Science to Management," pp. 451-457. Olsen and Olsen, Fredensborg. Ragueneau, O., Lancelot, C , Egorov, V., Vervlinmieren, J., Daoud, N., Rousseau, V., Popovitchev, v., Cauwet, G., and Deliat, G. (2(X)2). Present day biogeochemical transformations of inorganic nutrients in the Danube-north western Black Sea mixing zone. Estuarine Coastal Shelf Set, in press. Raymond, P. A., and Bauer, J. E. (2002). Bacterial consumption of DOC during transport through a temporate estuary. A^war. Microb. Ecol 22, 1-12. Richey, J. E. (1991). The biogeochemistry of a major river system: the Amazon case study. In "Biogeochemistry of Major World Rivers" (E. Degens, S. Kempe, and J. E. Richey, Eds.), pp. 57-74. Wiley, Chichester. Safmllah, S., Mofizuddin, M., Iqbal Ah, S. M., and Enalul Kabir, S. (1987). Biogeochemical cycles of carbon in the rivers of Bangladesh. In "Transport of Carbon and Minerals in Major World Rivers" (E. T. Degens, S. Kempe, and W. Gan, Eds.), part 4, pp. 435-442. Mitteilungen Der GeologischPalaontologischen Institutes der Universitat Hamburg. Saliot, A., Cauwet, G., Cahet, G., Mazaudier, D., and Daumas, R. (1996). Microbial activities in the Lena River delta and Laptev Sea. Mar. Chem. 53, 247-254. Seitzinger, S. P., and Sanders, R. W. (1997). Contribution of dissolved organic nitrogen from rivers to estuarine eutrophication. Mar. Ecol. Prog. Sen, 159, 1-12. Sempere, R., and Cauwet, G. (1995). Occurrence of organic colloids in the stratified estuary of the Krka River (Croatia). Estuarine Coastal Shelf Sci. 40,105-114. Sempere, R., Charriere, B., Van Wambeke, F , and Cauwet, G. (20(X)). Carbon inputs of the Rhone River to the Mediterranean sea: Biogeochemical implications. Global Biogeochem. Cycles 14,669-681. Serratore, P., Rinaldi, A., Montanari, G., Ghetti, A., Ferrari, C. R., and VoUenweider, R. A. (1995). Some observation about bacterial presence in seawater related to mucilaginous aggregates in the Northern Adriatic Sea. Sci. Total Environ. 165, 185-192. Sharp, J. H., Cifuentes, L. A., Coffin, R. B., Pennock, J. R., and Wong, K. C. (1986). The influence of river variability on the circulation, chemistry and microbiology of the Delaware Estuary. Estuaries 9, 261-269. Sharp, J. H., Pennock, J. R., Church, T. M., Tramontano, J. M., and Cifuentes, L. A. (1984). The estuarine interaction of nutrients, organics and metals: A case study in the Delaware Estuary. In "The Estuary as a Filter" (V. S. Kennedy, Ed.), pp. 241-258. Academic Press, Orlando. Sholkovitz, E. R. (1976). Rocculation of dissolved organic and inorganic matter during the mixing of river water and seawater. Geochim. Cosmochim. Acta 40, 831-845. Sholkovitz, E. R., Boyle, E. A., and Price, N. B. (1978). The removal of dissolved humic acids and iron during estuarine mixing. Earth Planet. Sci. Lett. 40, 130-136. S0ndergaard, M., WiUiams, P. J. 1., Cauwet, G., Riemann, B., Robinson, C , Terzic, S., Woodward, M., and Worm, J. (2000). Organic carbon partitionning and DOC accumulation in marine plankton conmiunities: Controls, flux and fate of DOC in a mesocosm experiment. Limnol. Oceanogr. 45, 1097-1111. Spitzy, A., and Leenheer, J. (1991). Dissolved organic carbon in rivers. In "Biogeochemistry of Major Worid Rivers" (E. Degens, S. Kempe, and J. E. Richey, Eds.), pp. 213-232. Wiley, Chichester. Stepanauskas, R., Leonardson, L., and Tranvik, L. J. (1999). Bioavailability of wetland-derived DON to freshwater and marine bacterioplankton. Limnol. Oceanogr 44,1477-1485. Telang, S. A., Pocklington, R., Naidu, A. S., Romankevitch, E. A., Gitelson, I. I., and Gladyshev, M. I. (1991). Carbon and minerals transport in major North American, Russian Arctic and Siberian Rivers: The St Lawrence, the Mackenzie, the Yukon, the arctic Alaskan Rivers, the arctic basin

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rivers in the Soviet Union and the Yenisei. In "Biogeochemistry of Major World Rivers, SCOPE Report 42" (E. T. Degens, S. Kempe, and J. E. Richey, Eds.), pp. 75-104. Wiley, Chichester. Thingstad, T. R, Hagstrom, A., and Rassoulzadegan, F. (1997). Accumulation of degradable DOC in surface waters: Is it caused by a malfunctioning microbial loop? Limnol Oceanogr. 42, 398-404. Thurman, E. M. (1985). "Organic Geochemistry of Natural Waters." Martinus Nijhoff/Dr W. Junk, Boston. Welker, C , and Nichetto, R (1996). The influence of mucous aggregates on the microphytobenthic community in the northern Adriatic Sea. RS.Z.N.I. Mar. Ecol 17,473^89. Wells, M. L. (2002). Marine colloids and metals. In "Biogeochemistry of Marine Dissolved Organic Matter" (D. A. HanseU and C. A. Carlson, Eds.), pp. 367^04. Academic Press, San Diego. Whitehouse, B. G., MacDonald, R. W, Iseki, K., Yunker, M. B., and McLaughlin, F. A. (1989). Organic carbon and colloids in the Mackenzie River and Beaufort Sea. Mar. Chem. 26, 371-378. Williams, P. J. L. (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. Mar. Chem. 51,17-29. WiUiams, P. M., and Druffel, E. R. M. (1987). Radiocarbon in dissolved organic matter in the central North Pacific Ocean. Nature 330,246-248. Zhang, S., Gan, W. B., and Ittekkot, V. (1992). Organic matter in large turbid rivers: The Huanghe and its estuary. Mar. Chem. 38, 53-68. Zweifel, U. L. (1999). Factors controlling accumulation of labile dissolved organic carbon in the Gulf of Riga. Estuarine 48,357-370. Zweifel, U. L., Wikner, J., Hagstroem, A., Lundberg, E., andNorrman, B. (1995). Dynamics of dissolved organic carbon in a coastal ecosystem. Limnol. Oceanogr 40, 299-305.

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Chapter 13

Sediment Pore Waters D a v i d J. Burdige Department of Ocean, Earth, and Atmospheric Sciences Old Dominion University Norfolk, Virginia I. Introduction A. Scientific Background II. Dissolved Organic Carbon (DOC) in Sediment Pore Waters A. General Observations B. An Advection/Diffusion/ Reaction Model for Sediment DOC Cycling Controls on DOC Concentrations with Depth in Surficial ("Shallow") Sediments D. Pore Water DOC Profiles in Deep Sediment Cores III. Dissolved Organic Nitrogen (DON) IV. Dissolved Organic Matter (DOM) Compositional Data A. Volatile Fatty Acids (VFAs) B. Amino Acids C. Carbohydrates V. The Role of Benthic DOM Fluxes in the Ocean Carbon and Nitrogen Cycles

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

A. Benthic DOC Fluxes B. Benthic DON Fluxes C. The Extent to Which Benthic DOM Fluxes Affect the Composition and Reactivity of Deep-Water DOM VI. The Role of Pore-Water DOM in Sediment Carbon Preservation VII. Conclusions and Suggestions for Future Research Acknowledgments Appendix: A Description of the DOM Advection/Diffusion/ Reaction Model Anoxic, Nonbioturbated Sediments (ANS Model) Bioturbated and/ or Bioirrigated Sediments (BBS model) References

611

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David J. Burdige

I. INTRODUCTION Dissolved organic matter (DOM) in marine sediment pore waters plays an important role in sediment carbon remineralization and may also be involved in sediment carbon preservation (see discussions in Krom and Westrich, 1981; Hedges, 1992, Henrichs, 1992, 1993; Hedges and Keil, 1995). It is also important in oceanic and sediment nutrient (N and P) cycling (e.g., Burdige and Zheng, 1998). Finally, DOM in sediment pore waters plays a role in pore water metal complexation and therefore affects dissolved metal and metal-complexing ligand fluxes from sediments (Elderfield, 1981; Skrabal et aL, 1997, 2000). Of the articles cited above that discuss the role of DOM in sediment carbon remineralization and preservation, only Krom and Westrich (1981) summarizes the known information on the concentrations and composition of pore water DOM (also see Thurman, 1985, for similar discussions). In spite of the 15- to 20-year gap since the presentation of such sunmiaries, the goal here is not, however, to simply update these inventories. Rather, I wish to specifically examine DOM in sediment pore waters with reference to several specific questions: 1. What do we know about the composition and reactivity of pore-water DOM, with particular reference to its role in sediment organic matter remineralization? 2. What do we know about the controls on pore-water DOM concentrations? 3. What is the role of benthic DOM fluxes in the global ocean cycles of carbon and nitrogen? 4. What is the role of pore-water DOM in sediment carbon preservation? The discussion in this chapter will focus on examining these questions in estuarine and continental margin sediments, in part because much of the detailed work in recent years on sediment DOM cycling has occurred in these sediments. Since the vast majority of carbon preservation and remineralization in all marine sediments occurs in estuarine and continental margin sediments (Hedges, 1992; Hedges and Keil, 1995; Middelburg etal, 1997), there is also important scientific justification for taking this approach. A.

SCIENTIFIC B A C K G R O U N D

To begin this discussion I start with the general observation that DOM in sediment pore waters is a heterogeneous collection of organic compounds, ranging in size from relatively large macromolecules (e.g., dissolved proteins or humic substances) to smaller molecules such as individual amino acids or short-chain organic acids (e.g., Krom and Westrich, 1981; Burdige and Martens, 1990; Henrichs, 1992). Concentrations of DOM in pore waters (both dissolved organic carbon [DOC] and dissolved organic nitrogen [DON]) are generally elevated over bottom-water

Sediment Pore Waters

613

values (up to an order of magnitude), implying that there is net production of this material in sediments as a result of remineralization processes. Much of the total pore water DOC and DON is of relatively low molecular weight ( 10) relative to the more N-rich DOM that escapes the sediments as a benthic flux (which has a C/N ratio of ~4-6 at site M3 and ^^2-4 at site S3; Burdige and Zheng, 1998). Similar trends in DOM elemental ratios have been observed in other sediments (Blackburn et aL, 1996; Landen-Hillmeyr, 1998), and were explained as being due to diffusional loss of low C/N ratio DOM produced during the initial hydrolysis of fresh (i.e., low C/N ratio) detrital organic matter near the sediment surface. This explanation is consistent with our Chesapeake Bay observations discussed above and discussions in Burdige and Gardner (1998), regarding the spatial separation in sediments between processes that produce the initial high-molecular-weight intermediates of sediment POM remineralization, and processes responsible for the production of refractory DOM in sediment pore waters (i.e., pLMW DOM). This spatial separation of HMW DOM and pLMW DOM net remineraUzation rates can also be inferred from the model results in Fig. 4. Based on these results the

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647

uncoupling between DOM that escapes as a benthic flux and that which accumulates in pore waters can be explained if one assumes that the C/N ratio of humic-like, pLMW DOM is greater (i.e., more carbon-rich) than that of this reactive HMW DOM, which appears to be a reasonable assumption (see Burdige, 2001, for further details). Results in Fig. 4 also show that in spite of the fact that model-derived concentrations of HMW DOC were significantly lower than those of pLMW DOC, their pore water gradients near the sediment surface were similar in magnitude. This observation is further quantified in Table I where it can be seen that calculated HMW DOC benthic fluxes are --50-80% of the total benthic DOC flux. Thus, in both anoxic and mixed redox sediments, model results suggest that fluxes of both refractory and reactive DOM can be similar in magnitude. In another approach to this problem, Bauer et al. (1995) examined the stable carbon and radiocarbon content of pore water DOM from the upper ~5 cm of two contrasting eastern North Pacific sediments (in Santa Monica Basin and at a 4100 m water depth site on the continental rise at the base of the Monterey deep-sea fan). At both sites the 8^^C values of the pore-water DOC had values consistent with a predominant marine source (—21 to —22 %o) and the pore-water DOC was greatly enriched in ^"^C as compared to bottom-water DOC (approx. —150 to —250 %oin the pore waters versus approx. —500 %o in the water column). While pore-water DOC profiles at both site predict diffusive, benthic DOC fluxes that are near-equal to the sediment carbon oxidation rate, this sediment DOC source does not appear to have a significant impact on bottom-water DOC radiocarbon values. To explain this latter observation, Bauer et al. (1995) suggested that either ^"^C-enriched pore water DOC is not released from the sediments in significant quantities (i.e., they have overestimated the benthic DOC flux in their simple benthic flux calculation) or that this radiocarbon-enriched, sediment-derived DOC "does not persist in the water colunm" (i.e., it is sufficiently reactive that it is rapidly remineralized). Consistent with the first suggestion is the fact that benthic DOC fluxes at these sites that I estimate using reported sediment carbon oxidation rates and Eq. [5] are seven to eight times smaller than those calculated by Bauer et al. (1995). At the same time though, the model results discussed earlier in this section (and shown in Fig. 4) suggest that benthic DOC fluxes may contain both refractory and reactive DOM components, which are likely to have different radiocarbon ages (e.g., see Eglinton et a/., 1997, for a discussion of analogous studies of radiocarbon ages of biomarker compounds in sediment POC). This would imply that the radiocarbon values determined by Bauer et al. (1995) could result from refractory (radiocarbon-depleted) and reactive (radiocarbon-enriched) components. This reactive sediment-derived DOC would likely undergo remineralization in the bottom waters, thus further minimizing the importance of sediment pore waters as a source or refractory DOC to the deep ocean.

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David J. Burdige

While the discussion here shows that benthic DOC fluxes may not be as important a source of refractory DOC to the water colunm as once suggested, recent results suggest another way that sediment pore-water DOC dynamics could play a role in explaining the ^"^C age of deep-water DOC. In recent studies, Guo and Santschi (2000) observed that simple desorption of colloidal (>1 kDa) organic matter from continental margin sediments yields DOC that is substantially older than the bulk sediment organic matter (^3000 years vs 700 years, respectively). Desorption of this material from sediments in continental margin benthic nepheloid layers, coupled with transport of this material to the deep ocean (e.g., Bauer and Druffel, 1998), could then play an important role in explaining the age of deepocean DOC. Assuming that this old, desorbed DOC is in some kind of reversible equilibrium with the sediment pore waters while in the sediments, the relatively refractory nature of the bulk pore water DOC pool (^pLMW DOC) would then aid in the aging process of this sorbed organic matter while it is resides in the sediments.

VI. THE ROLE OF PORE-WATER DOM IN SEDIMENT CARBON PRESERVATION As was discussed at the beginning of this chapter, two common models for sediment carbon preservation, the geopolymerization model (Nissenbaum et ai, 1972; Tissot and Welte, 1978; Krom and Westrich, 1981) and the mesopore protection model (Mayer, 1994a,b; Hedges et ai, 1999), suggest that DOM in sediment pore waters may play an important role in the preservation process. A third model for carbon preservation, the selective preservation of refractory biomacromolecules (Hatcher and Spiker, 1988; de Leeuw and Largeau, 1993) is generally thought to not involve DOM intermediates. In its simplest sense, the geopolymerization model involves condensation reactions in which low-molecular-weight dissolved organic compounds (such as amino acids or simple sugars; i.e., mLMW DOM compounds in the PWSR model) react to form higher-molecular-weight dissolved humic substances. The continued condensation of these dissolved humics is thought to eventually lead to the formation of particulate material such as humin or kerogen, and to the preservation of this organic matter in sediments. Although there have been numerous studies of these condensation reactions in the lab, there is little direct evidence for their occurrence in nature (Hedges, 1988; Henrichs, 1992). Furthermore, based on our molecularweight data (Burdige and Gardner, 1998), we noted that if these reactions do occur on early diagenetic time scales their products still have relatively low molecular weights (i.e., less than 3 kDa). In the mesopore protection model, DOM sorption is proposed to occur in small mesopores on mineral surfaces, where the sorbed DOM is physically protected

Sediment Pore Waters

649

from attack by microbial enzymes (Mayer, 1994a,b, 1999; Hedges and Keil, 1995). The mesopore protection model, however, is not mutually exclusive of the occurrence of geopolymerization reactions, since rates of abiotic condensation reactions may be accelerated in mesopore sites, by either steric- or concentration-related phenomena, thus further enhancing the preservation of sorbed DOM (Mayer, 1994b; Collins et al, 1995; Hedges and Keil, 1995). This latter point is of some interest, since under most natural conditions aqueous phase (i.e., pore water) geopolymerization reactions are thought to be quite slow (Hedges, 1988; Alperin et al, 1994). Since mesopore size places some constraint on the upper limit of the size/molecular weight of DOM molecules that can be taken up in mesopores (Mayer, 1994b), the relatively small size ( 10 x 10^ km^) with variable vegetation and soil conditions. Consequently, DOC concentrations vary significantly between rivers (Table I). There are also significant seasonal differences in river TOC concentration, as reported for the Lena River by Cauwet and Sidorov (1996). The maximum concentration (980 ^M) was found during the maximum water discharge in early summer, followed by a lower concentration (700 /xM) in the summer and autumn, and the lowest concentration (310 /xM) during winter. The mean annual, discharge-weighted, concentration was 830 jiM. It should be noted that several investigations were performed after maximum water discharge in summer, and not always is the date of sampling given in the literature. To get average discharge weighted concentrations, DOC concentrations were multiplied by each river discharge and divided by the total annual discharge (Table II). Another uncertainty is that around one-third of the discharge to the Arctic Ocean takes place through smaller rivers and creeks that are not included in Table I. An alternative approach for estimating an average discharge weighted concentration for the rivers entering a given area is to sample the estuary and the surrounding sea and make a DOC versus salinity plot (Fig. 2). Assuming that DOC behaves conservatively, the intercept at 5 = 0 corresponds to a discharge weighted mean of the rivers entering the area investigated. This estimate includes the seasonal variability, as the residence time of the runoff on the Eurasian shelves has been estimated to ^ 3 years (Schlosser et al, 1994). This approach is not suitable for the Beaufort Sea area, where the Mackenzie River discharges, as the residence time of the surface water on the shelf in summer is short (Macdonald et al, 1989). The Mackenzie River dominates the discharge from North America into the Arctic Ocean and it is thus more straightforward to evaluate the DOC concentration in the runoff from this continent, than from the Eurasian. The regression lines of Figs. 2C and 2D (from the Laptev Sea region) are in excellent agreement with an intercept of 579 and 580 /xM. The data of Fig. 2B fall

669

DOC in the Arctic Ocean Table I Reported Concentrations of DOC and TOC in Arctic Rivers No.

River

1 2 3 4 5 6 7 8 9 10 11 B.S.

Pechora Ob Pyr Yenisey Katanga Olenek Lena Yana Indigirka Kolyma Mackenzie Yukon

DOC {^iM)

111 850 538 to 558 232 to 264 404 387 375 to 863 357 to 733

TOC (^lM)

References

Shelf seas

1083 592 to 733 558 617 525 600 792 to 842 558 to 611 642 to 754 389 to 675 642 to 1050 476 to 833

d d,g d d,e,g d d,e a, c, d, e, g c, d, e, g c,d,g c,d,g b,dj,g

Barents Kara Kara Kara Laptev Laptev Laptev Laptev East Siberian East Siberian Beaufort Bering

b,f,g

Note. The Yukon river enters the Bering Sea (B.S.), outside the range of Figure 1, but most of its w^ater enters the Arctic Ocean through Bering Strait. "Cauvet and Sidorov (1996). ^Degens et al. (1991). ^Fitznar (1999). '^Gordeev et al (1996). ^Lobbesetfl/. (2000). ^Pocklington (1987). ^Telang et al (1991).

Table II DOC Flux from Major Rivers into the Arctic Ocean (Lobbes et al, 2000) River Mezen

Ob Yenisey delta Olenek Lena delta Yana Indigirka Kolyma Mackenzie Total

Discharge (km"^ year~^)

Drainage basin area (km^ 10^)

DOC flux D O C (jLtM)

(lO^gCyear-i)

56

1006

2,990 2,440

1,805

735 711 850 538 232 404 387 640

248 3,690 4,860 323 3,380 85 241 458 1,917

10,994

636^^

15,200

21 419 569 32 524 31 50 98 249

2,430

1,993

198 244 305 526

"The mean concentration is computed as (DOC flux)/(discharge) x (12).

CD

o o

8 8 CO

(£>

o o

8

in

o o

"3000 /xM) was measured in the bulk of the bottom ice on May 14 (Smith et al, 1997). However, the volume with this high concentration is small and thus the integrated contribution of DOC from ice to the underlying water mass is small. Measurements of DOC release rates by ice algae were performed by Gosselin et al (1997) along the track of the Arctic Ocean section in 1994. The release rate varied from less than 25 /xmol m~^ day~^ to 1600 lb 1500 /xmol m~^ day~^ with the highest rates in the Chukchi Sea. Thomas et al (1995) collected three ice cores of more than 2 m length in the Fram Strait. In two of these the DOC concentration was mostly below 100 /xM all through the core. In the third the concentration was close to 100 /xM in the top ^1.8 m, and increased to a maximum of ~700 /xM some 10 cm from the bottom. This increase was explained by a combination of DOM excretion by biota and decomposition of organisms (Thomas et al, 1995). The mean bulk concentration of DOC in sea ice from the central Arctic Ocean is 316 ± 99 /xM (Melnikov, 1997). If 1 m of ice melts annually, the concentration in the top 50 m (typical winter surface mixed layer (Rudels et al, 1996)) would increase by just over 6 /xM. A further source of DOC from biological processes is release from the sediment surface caused by decomposition of particulate organic material. Hulth et al (1996) measured DOC concentrations in the range of 500 to 8000 /xM in pore

674

LeifG. Anderson

water in the Svalbard area. The lowest concentrations were found at stations east of Svalbard, where also a significant inverse linear correlation (r^ = 0.849) of DOC concentrations with a sediment reactivity index (defined as sediment oxygen consumption rate normalized to the organic content) was found. This suggests a coupling between reactivity of organic matter in sediment and DOC lability in pore water. In a study of the eastern Eurasian Basin and adjacent shelves (Hulthe and Hall, 1997), DOC fluxes out of the sediment were evaluated to be in the range from close to zero to 3.6 mmol m"-^ day~^. The highest fluxes were found on the shelves and the lowest in the deep basins and on the slopes. A positive correlation of the DOC and dissolved inorganic carbon fluxes was observed, with DOC constituting up to 50% of the total benthic carbon flux at stations with the highest total benthic carbon fluxes. This indicates that the fraction of DOC that is oxidized to inorganic carbon is decreasing with increasing decomposition rates.

III. COMPOSITION AND DISTRIBUTION OF DOC WITHIN THE ARCTIC OCEAN Before the transport of DOC to and from of the Arctic Ocean is discussed, the quality of the terrigenous DOM has to be considered. Does it flow with the water as a biogeochemically stable solute or is it available to diagenetic alteration or photochemical decomposition? Several investigations have studied the composition of DOM in rivers entering the Arctic Ocean (e.g., Gordeev et al, 1996; Cauwet and Sidorov, 1996; Lara et al, 1998; Lobbes et ai, 2000) as well as in the Arctic Ocean itself (e.g., Wheeler ^r a/., l991;OpsahletaL, 1999; Kattner^r^/., 1999). One general conclusion is the stability of terrigenous DOC in the surface waters of the Arctic Ocean. Except for the Unear mixing line of runoff and seawater in a DOC vs salinity plot, the fairly constant composition of the DOM in all of the Arctic Ocean supports this conclusion.

A. LiGNiN OXIDATION PRODUCTS AND STABLE CARBON ISOTOPES The most useful quantitative tracers of terrestrial organic matter are lignin oxidation products, which have been determined in runoff to the Arctic Ocean (Opsahl et ai, 1999; Lobbes et al, 2000) and in the surface waters of the Arctic Ocean (Opsahl et al, 1999; Kattner et al, 1999). Kattner et al (1999) determined lignin in the "humic" fraction of DOM and used this as a tracer for terrigenous influence, with the result that the riverine-derived freshwater contribution to the Laptev Sea is 8 to 30%. Combining this proportion with DOC concentrations in the Lena River and Laptev Sea indicates that about 60% of the DOC in the surface layer of the Laptev Sea and adjacent Eurasian Basin would be of terrigenous

675

DOC in the Arctic Ocean

origin. In contrast, terrigenous dissolved organic nitrogen (DON) only accounted for 20 to 30% of the total DON (Kattner et al, 1999). However, as stressed by the authors, the distribution of DON is generally more influenced by biological processes, making this last estimate more uncertain. The fraction of terrigenous DOM in surface waters of the central Arctic Ocean was estimated from the carbon-normalized yields of lignin oxidation products (Ae) and (5^^C in ultrafiltered dissolved organic matter (UDOM) (Opsahl et al, 1999), resulting in 5-22% and 16-33%, respectively. The UDOM represents the highmolecular-weight fraction of DOM (>1 kDa), which is about 20-30% of total DOM. In Fig. 3 the mean values (ibvariability) of samples from the Kara Sea (low (5^^C), the polar surface water (medium 5^^C), and deep Fram Strait and Greenland Sea (high 34.5, which likely is a result of these samples being deep waters and thus have a signal affected by decay of sinking organic particulate matter. A large variability in the C/N ratio of DOM, ranging from 10 to 40 with a peak around 15, was also observed in the Fram Strait (Lara et al, 1998). The variable C/N ratio in marine dominated waters is a result of variable DON concentrations. This makes the C/N ratio less useful for

677

DOC in the Arctic Ocean

30-

C/N = -0.943*S + 49.7 F^= 0.2799

25 H

,

O Q

o o o

I 20. c3

X

o

o o

15H

X X

o^o

X

10-

I

28

I

29

30

—r32 Salinity

31

33

"~T—

34

35

Figure 4 C/N ratio in dissolved organic matter (DOC/DON) in the eastern Eurasian Basin. Open circles represent S 34.5. The linear regression line is fitted to the open circles. Data are from Fitznar (1999).

quantitative computations, but it is valuable as a qualitative tracer of terrigenous DOM.

C. DISTRIBUTION Too few DOC data are available from the central Arctic Ocean surface waters to produce a map of the concentration distribution. Several processes considerably influence the distribution by producing and consuming DOC. The relative importance of these processes can be seen in a DOC versus salinity plot of the surface waters with 590 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

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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

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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

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Pacific/Indian Sector

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32' 80°N

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40°S

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av^ace NPOH^//

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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).

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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),

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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

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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

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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

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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 (