Long-term Ecological Change in the Northern Gulf of Alaska

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LONG-TERM ECOLOGICAL CHANGE IN THE NORTHERN GULF OF ALASKA

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LONG-TERM ECOLOGICAL CHANGE IN THE NORTHERN GULF OF ALASKA Edited by Robert B. Spies

Elsevier Radarweg 29, PO Box 211, 1000 AE Amsterdam, The Netherlands The Boulevard, Langford Lane, Kidlington, Oxford OX5 1GB, UK

First edition 2007 Copyright © 2007 Elsevier B.V. All rights reserved No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or by any means electronic, mechanical, photocopying, recording or otherwise without the prior written permission of the publisher Permissions may be sought directly from Elsevier’s Science & Technology Rights Department in Oxford, UK: phone (+44) (0) 1865 843830; fax (+44) (0) 1865 853333; email: [email protected]. Alternatively you can submit your request online by visiting the Elsevier web site at http://elsevier.com/locate/permissions, and selecting Obtaining permission to use Elsevier material Notice No responsibility is assumed by the publisher for any injury and/or damage to persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein. Because of rapid advances in the medical sciences, in particular, independent verification of diagnoses and drug dosages should be made Library of Congress Cataloging-in-Publication Data A catalog record for this book is available from the Library of Congress British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library ISBN-13: 978-0-444-52960-2 ISBN-10: 0-444-52960-8

For information on all Elsevier publications visit our website at books.elsevier.com

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We dedicate this book to the people of Alaska. We hope that it will help inspire both further support for research and wise human action in the Gulf of Alaska.


Acknowledgements

This book was made possible through the passionate work of government, academic and private scientists in the Northern Gulf of Alaska. Many of these hard-working scientists have dedicated their careers to some aspect of this productive ecosystem. Special thanks are due to the Exxon Valdez Trustee Council of 2000, who provided the funds for this project. This and previous councils recognized the wisdom of dedicating some small part of the trust funds to summarizing what we have learned since the 1989 oil spill about the ecosystem in which it occurred. Their vision for long-term monitoring and research is a great example for future political leaders. Graphic material for this book was prepared by Peter Veres. We are grateful to the many naturalist photographers who permitted use of their work.


Contributors

Kevin M. Bailey, Ph.D. Alaska Fisheries Science Center NOAA/NMFS 7600 Sand Point Way N.E., Building 4 Seattle, WA 98115 [email protected] Z. Morgan Benowitz-Fredericks, Ph.D. Department of Biology PO Box 351800 University of Washington Seattle, WA 98195 [email protected]. James L. Bodkin, Ph.D. Alaska Biological Science Center USGS 1011 E. Tudor Blvd. Anchorage, AK 99503 Wiebke J. Boeing, Ph.D, Department of Fishery and Wildlife Sciences 2980 South Espina, 132 Knox Hall PO Box 30003, Campus Box 4901 New Mexico State University Las Cruces, NM 88003-8003 [email protected]

Evelyn Brown, Ph.D. School of Fisheries and Ocean Sciences University of Alaska Fairbanks Fairbanks, AK 99775 [email protected] Theodore Cooney, Ph.D. Emeritus Professor School of Fisheries and Ocean Sciences University of Alaska Fairbanks, AK 99775-7220 PO Box 486 Choteau. Montana [email protected] Sara J. Iverson, Ph.D. Department of Biology Dalhousie University 1355 Oxford Street Halifax, NS Canada B3H 4J1 [email protected] Gordon H. Kruse, Ph.D. President’s Professor of Fisheries University of Alaska, Fairbanks Juneau Center for Fisheries & Ocean Sciences 11120 Glacier Highway Juneau, AK 99801-8677 [email protected]

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Contributors

John F. Piatt, Ph.D. Alaska Biological Science Center USGS Anchorage, Alaska Canada Fairbanks, AK Mailing address: Marrowstone Marine Station 616 Marrowstone Point Road, Nordland, WA 98358-9633 [email protected] Paul Reno, Ph.D. Department of Microbiology Oregon State University Coastal Oregon Marine Experiment Station Newport, OR Mailing address: Hatfield Marine Science Center 2030 SE Marine Science Drive Newport, OR 97365 [email protected] Stanley D. Rice, Ph.D. NOAA/NMFS Auke Bay Laboratory 11305 Glacier Highway Juneau, AK 99801-8626 [email protected]

Robert B. Spies PO Box 315 45100 Peterson St. Little River, CA 95456 [email protected] Alan M. Springer, Ph.D. Institute of Marine Sciences University of Alaska 1708 Marmot Hill Road Fairbanks, AK 99709 [email protected] Thomas Weingartner, Ph.D. Institute of Marine Sciences University of Alaska Fairbanks, AK 99775 [email protected]

Contents

1 Introduction

1

Robert B. Spies and Theodore Cooney

1.1. Why We Have Written This Book 7 1.2. Who Is the Audience for This Book, and What Is Its Scope? 8 1.3. Organization 8 References 9 2 Ecosystem Structure 11 2.1. Introduction 11 Robert B. Spies and Alan M. Springer

2.2. The Physical Environment of the Gulf of Alaska 12 Thomas Weingartner

2.2.1. Geomorphology 13 Local Atmospheric Forcing 17 Precipitation and Runoff 18 Winds 21 Heat Fluxes 24 2.2.2. Physical Oceanography 26 The Seasonal Cycle of Water Properties 26 Circulation over the Gulf of Alaska Shelf and Slope 30 2.2.3. Tides 41 2.2.4. Gulf of Alaska Basin 42 2.2.5. North Pacific Ocean 44 2.3. The Marine Production Cycle 47 Theodore Cooney

2.3.1. Introduction 47 2.3.2. The Annual Cycle of Production 48 Protected Inner Waters 55 The Open Ocean 57 Open Coastal and Shelf Waters 59

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Contents

2.4. The Transfer of Matter and Energy Through the Food Web 60 Theodore Cooney

2.4.1. Food-Web Structure – The Principal Forage Stocks 62 2.4.2. Efficient Foraging on Patches 70 2.4.3. Food-Web Complexity and Efficiency 71 2.5. Strategies for Survival 74 2.5.1. Introduction 74 Alan M. Springer

2.5.2. Introduction to Fishes

75

Theodore Cooney

2.5.3. Pink Salmon

76

Theodore Cooney

2.5.4. Pacific Herring

81

Theodore Cooney

2.5.5. Walleye Pollock

85

Kevin M. Bailey and Lorenzo Ciannelli

Introduction 85 Adaptations for Survival 87 Effect of Ecosystem Structure on Pollock Survival 91 Conclusions 92 2.5.6. Comparing Fish Life Histories 93 Theodor Cooney

2.5.7. Seabirds 94 Morgan Benowitz-Fredericks, A.S. Kitaysky and Alan M. Springer

Introduction 94 Seabirds in the Breeding Season 95 Focal Species 95 Phylogeny 96 Foraging Ecology and Reproductive Strategies 96 Foraging Ecology 97 Foraging Range and Habitat Use 100 Diet 101 Response to Changes in Prey 102 Life History Traits and Reproduction 103 Chick Strategies 104 Parental Behavior and Consequences for Survival 108 Other Factors Affecting Survival: Nesting and Predation 111 Summary 112 2.5.8. Marine Mammals 114 Sara J. Iverson, Alan M. Springer, and James Bodkin

Introduction 114 Offspring Constraints and Strategies 121

Contents

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Trade-offs and Consequences: Energy and Population Considerations 123 Predator Avoidance 130 Summary 131 2.5.9. Crabs and Shrimps 135 Gordon H. Kruse

Introduction 135 Distribution and Habitats for Survival 135 Feeding Strategies 137 Predator Defense Strategies 138 Reproductive Strategies 140 Larval Survival: Strategies of Timing (or “Match the Hatch”) 143 Larval Survival: Strategies of Space (or “Location, Location, Location”) 144 Implications of Survival Strategies 145 References 145 Mini-Glossary 169 3 Agents of Ecosystem Change 171 3.1. Introduction 171 Robert B. Spies

3.2. Climate 171 Thomas Weingartner

3.2.1. Introduction 171 3.2.2. Climate Forcing 172 3.2.3. Physical Environmental Variability in the North Pacific Ocean 176 El Niño Southern Oscillation (ENSO) 176 The Pacific Decadal Oscillation (PDO) 178 3.3. Geophysical Mechanisms 180 Robert B. Spies

3.3.1. Introduction 180 3.3.2. Tectonics and Earthquakes 180 3.3.3. Sediment Slumping 183 3.3.4. Volcanism 183 3.3.5. Tsunamis 186 3.3.6. Glaciers 187 3.4. Species Interactions 187 Gordon H. Kruse

3.4.1. Introduction 187 3.5. Marine Mammal Harvest and Fishing 192 Gordon H. Kruse and Alan M. Springer

3.5.1. Introduction 192

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Contents

3.5.2. Marine Mammal Harvests and Persecution 192 Great Whales 192 Pinnipeds 196 Sea Otters 199 3.5.3. Fisheries of the Northern Gulf of Alaska 201 Introduction 201 3.5.4. History of Fishing and Fishery Management in Alaska 203 Salmon 203 Herring 205 Groundfish 207 Shellfish 208 3.5.5. Direct and Indirect Effects of Fishing 211 3.5.6. Effects of Hunting and Fishing: Some Conclusions 218 3.6. Disease 219 Paul Reno

3.6.1. 3.6.2. 3.6.3. 3.6.4.

Introduction 219 Agents of Infectious Disease and their Lifestyles 220 Resistance to Infection and Disease 224 The Diversity of Infectious Agents Causing Diseases in Marine Animals 226 Diseases of Coral (http://www.coral.noaa.gov/coral_ disease/) 230 Diseases of Molluscs (http://www.pac.dfo-mpo.gc.ca/sci/shelldis/ title_e.htm) 230 Diseases of Crustaceans (http://www.pac.dfo-mpo.gc.ca/sci/shelldis/ pages/hematcb_e.htm) 233 Diseases of Teleosts (http://wfrc.usgs.gov/capfishhealth.htm) 233 Diseases of Reptiles (http://nationalzoo.si.edu/ConservationAndScience/ AquaticEcosystems/SeaTurtles/deem.cfm) 236 Diseases of Birds (http://www.nwhc.usgs.gov/pub_metadata/ index.html) 236 Diseases of Marine Mammals (http://www.nmfs.noaa.gov/ pr/health/) 237 3.6.5. Population Effects and Implications 238 3.7. Contaminants 241 Robert B. Spies and Stanley Rice

3.7.1. Significance of Life Stage at Time of Impact 246 3.7.2. Floating Oil and Surface-dwelling Animals 246 References 247 Mini-Glossary 256

Contents

4 Long-Term Change 259 4.1. Introduction 259 Robert B. Spies and Thomas Weingartner

4.2. Atmosphere and Ocean 265 Thomas Weingartner

4.2.1. Introduction 265 4.2.2. Gulf of Alaska Shelf 266 Winds 266 Variability in Surface Heating and Cooling and Shelf Temperatures 266 Variability in Runoff and Shelf Salinity 269 4.2.3. Gulf of Alaska Basin 273 4.3. Zooplankton 274 Theodore Cooney

4.4. History and Production Trends in Salmon 278 Theodore Cooney

4.4.1. 4.4.2. 4.4.3. 4.4.4. 4.4.5.

Introduction 278 Resource Use and Management 278 Broadscale Trends in Time and Space 280 Prince William Sound: A Pink Salmon Case History 284 What’s Behind the Large-Scale Patterns in Salmon Catch and Production? 287 4.5. Pacific Herring 290 Evelyn Brown

4.5.1. Resource Use and Management 290 4.5.2. Effects of Climate on Gulf of Alaska Herring 293 4.6. Groundfish 300 Wiebke J. Boeing, Michael H. Martin, and Janet T. Duffy-Anderson

4.6.1. 4.6.2. 4.6.3. 4.6.4. 4.6.5. 4.6.6.

Introduction 300 Walleye Pollock 303 Pacific Cod 303 Arrowtooth Flounder 305 Pacific Ocean Perch 305 Explaining Population Change 306 Climate forcing 306 Biological controls 308 Fisheries effects 309 Summary 310 4.7. Seabirds in the Gulf of Alaska 311 Alan M. Springer

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Contents

4.7.1. 4.7.2. 4.7.3. 4.7.4. 4.7.5.

Introduction 311 Long-term Changes in Common Murres 313 Long-term Changes in Tufted Puffins 320 Long-term Changes in Black-legged Kittiwakes 321 Long-term Changes in Other Species 323 Storm Petrels 323 Cormorants 323 Murrelets 324 Miscellaneous Species: PWS 324 4.7.6. Causes of Long-term Change in Seabirds 325 Changing Abundance 326 4.7.7. Changing Productivity 329 4.7.8. Conclusions 332 4.7.9. Summary 334 4.8. Population Ecology of Seabirds in Cook Inlet 335 John F. Piatt and Ann M.A. Harding

4.8.1. Introduction 335 4.8.2. The Cook Inlet Ecosystem 337 4.8.3. Response of Seabirds to Variability in Prey 339 Form of Response 339 Variability in Response 340 4.8.4. Population Dynamics of Seabirds in Cook Inlet 344 Population Parameter Indices 347 4.8.5. Long-term Changes in the Gulf of Alaska Marine Environment 350 4.9. Marine Mammal Populations 352 Alan M. Springer, Sara J. Iverson, and James L. Bodkin

4.9.1. Harbor Seals 352 4.9.2. Steller Sea Lions 355 4.9.3. Sea Otters 357 4.9.4. Potential Causes of Marine Mammal Population Change 361 4.9.5. Harbor Seal Decline 363 4.9.6. Steller Sea Lion Decline 368 4.9.7. Sea Otter Population Changes 374 4.9.8. Conclusions 375 4.10. Crabs and Shrimps 378 Gordon H. Kruse

4.10.1. Introduction 378 4.10.2. Long-term Dynamics of Shrimp Stocks 379 4.10.3. Long-term Dynamics of Crab Stocks 380

Contents

4.10.4. Climate Forcing 383 Is Climate Important to Invertebrate Populations? 383 Shrimp and the Match–Mismatch Hypothesis 385 Red King Crab and the Match–Mismatch Hypothesis 386 Reconciling the Match–Mismatch Hypothesis among Shrimp and Crabs 387 Other Potential Climate-forcing Mechanisms and Red King Crab 388 Climate Forcing and Tanner Crabs 388 4.10.5. Biological Controls 389 4.10.6. Fishing Effects 392 4.10.7. Conclusions 393 References 394 Mini-Glossary 418 5 The Exxon Valdez Oil Spill 419 Stanley D. Rice, Jeffrey W. Short, Mark G. Carls, Adam Moles and Robert B. Spies

5.1. 5.2. 5.3. 5.4.

Introduction 419 Pre-spill Conditions 422 History of the Spill 423 Oil Fate: Transport, Weathering, and Persistence 428 5.4.1. Contaminants in Prince William Sound Prior to the Exxon Valdez Oil Spill 432 5.4.2. Initial Fate of the Oil 433 5.4.3. Oil Cleanup Efforts 437 5.4.4. Long-term Oil Persistence 440 Early Surveys (1989–1993) 440 Post-1993 Surveys 442 5.4.5. Weathering and Bioavailability of Persistent Oil 445 5.5. Effects of the Spill on Aquatic Organisms 447 5.5.1. Acute Effects 449 Birds 450 Sea Otters 450 Seals 451 Sea Lions 452 Whales 452 5.5.2. Short-term Effects of the Spill 454 Pacific Herring 455 Pink Salmon Juveniles 457 Intertidal and Subtidal Communities 461 5.5.3. Long-term Effects of the Spill 466 Pink Salmon Embryos 466

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Contents

Mussels 477 Sea Otters and Harlequin Ducks 480 5.6. Indirect Interactions 486 5.7. Expectations, Certainty, and Final Lessons from the Exxon Valdez Oil Spill 487 5.8. The Legacy of the Exxon Valdez Oil Spill 489 5.9. Appendices 491 Appendix A 491 Appendix B 503 Properties and Composition of Fresh and Weathered Alaska North Slope Crude Oil 503 References 506 Mini-Glossary 520 6 Long-Term Changes in the GOA: Properties and Causes 521 Robert B. Spies, Theodore Cooney, Alan M. Springer, Thomas Weingartner and Gordon H. Kruse

6.1. Introduction 521 6.2. Forces of Change 528 6.2.1. Climate 528 6.2.2. The Exxon Valdez Oil Spill 531 6.2.3. The 1964 Earthquake 534 6.2.4. Effects of Harvesting 535 6.3. Characteristics of Ecosystem Change 535 6.3.1. Understanding Productivity is Key 536 6.3.2. Responses to Climate Changes Vary at the Mesoscale 537 6.3.3. Animals with High Reproductive Rates Periodically Dominate 538 6.3.4. Top-down Forces Have been Underestimated 538 6.3.5. Bottom-up and Top-down Operate Together 543 6.3.6. Contaminants and Disease as Predators 548 6.3.7. Structure of the Ecosystem Affects Ecological Change 549 6.3.8. Interaction of Multiple Forcing Factors and Cascading Effects 549 6.3.9. Prediction of Ecological Change 550 References 555 Index 561

Chapter 1

Introduction Robert B. Spies and Theodore Cooney

Our species has had a long and complex relationship with the sea. It enabled us to reach all of earth’s continents during our ancient migrations, for modern commerce, and for warfare. Its fish, shellfish, algae, and mammals supply much of our food, and today, new pharmaceuticals. Its mineral resources are used for producing energy and modern manufactured goods, and its huge chemical engine replenishes atmospheric oxygen. Sadly, the sea has also become the most convenient repository for the complex waste streams of our modern economy. Through millions of years of our evolution on a planet with three-quarters of its surface covered by water, we have developed deep, enduring, practical and spiritual ties with the sea. Only in the past 50 years have we really begun to understand that the cumulative actions of billions of humans are changing the sea on a large scale. Yet, in the absence of a clear understanding of natural changes in oceanic waters, including cycles in temperature, salt content, current patterns, and plant and animal populations, we cannot determine how exactly we are affecting the sea. Today, global climate change threatens us through sea level rise, increasingly violent storms, new patterns of precipitation, and alteration of fish populations. A majority of the earth’s exploited fish stocks are in decline; polar ecosystems are contaminated with exotic persistent chemicals that could disrupt human and wildlife reproduction. Coastal habitats and their spiritual values are under siege, as much of the earth’s burgeoning population is located near the sea. Therefore, we need, more than ever, a deeper understanding of how the ocean works, so that we can limit our impacts. Fishermen have long known that the harvestable species populations are always changing. Since the dawn of modern science in the 1700s, we have studied the ocean to see why this is so. Progress has been frustratingly slow. We are generally unable to predict changes from first principles, even after several centuries of conceptualizing and grappling with an increasing awareness of the complexity of the sea. However, a few key insights are beginning to emerge. Our understanding of the marine ecosystem is probably at a similar level as our understanding of astronomy was in 1920. Nowhere is ecosystem change more profoundly apparent than in the Gulf of Alaska. The rise and fall in the populations of salmon in the twentieth century, the decline of commercial shrimp and crab populations, the near extinction of sea otters, and other changes all attest to the frequent, large-scale, fundamental shifts in the Long-Term Ecological Change in the Northern Gulf of Alaska Robert B. Spies (Editor) © 2007 Elsevier B.V. All rights reserved.

2

Long-Term Ecological Change in the Northern Gulf of Alaska

Gulf’s animal populations (see Box 1.1). Figures 1(A–G) depict examples of these dramatic changes, some of which last for centuries. There are powerful natural and anthropogenic forces that drive changes in the northern Gulf of Alaska. With regard to natural forces, the northern Gulf of Alaska has strong seasonal, annual, interannual, decadal, and longer-term cycles. The atmosphere and ocean are intimately linked. Meteorological changes find expression in the ocean and in coastal watersheds. In fact, the entire North Pacific Ocean appears to fluctuate broadly from east to west in quasi-long-term cycles, ranging from a few years to decades or even longer. The strength and geographic position of the Aleutian Low Pressure system, the dominant meteorological feature of the North Pacific Ocean, forces many of the ocean’s responses to climate change – on a range of time and space scales. In addition, the North Pacific Ocean warms during El Ni~ nos every 4 to 5 years, but the relationship to the Aleutian Low Pressure is not well understood. Contemporary studies have demonstrated that changes in the ocean can occur in as little as a day after atmospheric influences, but oceanic conditions also track the climate over extended periods – decades and even longer. In the Gulf of Alaska, single storms, abnormal seasons, the inter-annual El Ni~ no–La Ni~ na cycle, and 20- to 50-year climate fluctuations leave a variety of historical records. A recently assembled history

BOX 1.1: CHANGING ANIMAL POPULATIONS IN THE GULF by Robert Spies and Theodore Cooney Animal populations in the Gulf of Alaska have waxed and waned for thousands of years, but the only historical evidence for such long-term change comes from the sockeye salmon, recently revealed in lake sediments on Kodiak Island (Fig. 1.1). During the nineteenth century, the sea otter and whale harvests, and in the twentieth century the harvest of a half million whales in the North Pacific Ocean (e.g., Fig. 1.2), had direct effects and may have set in motion, cascading long-term changes in the Gulf ecosystem. In the twentieth century, several rearrangements of the Gulf ecosystem occurred, favoring different species in different decades. The twentieth century changes were first evident in fluctuations in the salmon fisheries (Fig. 1.3), and then the commercial fin and shellfish fisheries in the 1950s to 1970s (Fig. 1.4). Most recently, systematic studies of crustacean and fish populations in the 1970s set the foundation for recording a documented shift in the Gulf, beginning at about the middle of that decade (Fig. 1.5). There were also dramatic declines in marine mammal populations (Fig. 1.6). The huge Exxon Valdez oil spill in 1989 had major impacts on intertidal communities, seabirds, and marine mammals in parts of the Gulf.

Introduction

10 9

δ15N

8 7 6 5 4 3 −250

0

250

500

750 1000 1250 1500 1750 2000 YEAR AD

Figure 1.1: Long-term record of sockeye salmon (Onchorhynchus nerka) abundance in the northern Gulf of Alaska. (after Finney et al., 2002). The data are nitrogen isotope ratios (δ15N) from the sediment cores of Akalura Lake, Kodiak Island, plotted by year, with the ratio value increasing in proportion to the population size.

Figure 1.2: Fin whale (Balenoptera physalus) harvest in the North Pacific Ocean, 1924–1987. Harvested biomass in metric tons. Each dot represents the cumulative total biomass caught in a 100-km2 area (after Springer et al., 2003).

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CATCH (millions of salmon)

80 70 60 50 40 30 20 10 0 1990

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Figure 1.3: Commercial salmon catches in Alaska during the twentieth century (after Francis and Hare, 1994).

350000 PACIFIC OCEAN PERCH CATCH (metric tons)

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Figure 1.4: Catch of Pacific ocean perch (Sebastes alutus) (after Hanselman et al., 2005).

Introduction

a

b

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Figure 1.5: Trawl catches in Pavlov Bay, Alaska Peninsula, show the changes in the Gulf’s inner-shelf communities from (a) crustacean dominance in the 1960s and 1970s, (b) in transition with a mixture of shrimp and fish, 1977–1980, and (c) finally, to flatfish and gadid (cod-like fishes) dominance after 1981. (Photographs, courtesy of Paul Anderson, ret. NOAA, National Marine Fisheries Service. Also see Piatt and Anderson, 1996).

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c b

a

Figure 1.6: The decline of the Steller sea lion (Eumetopias jubatus) stocks from the late 1970s through the 1980s. Photographs of the same beach at Ugamak Island taken in (a) 1969, (b) 1979, and (c) 1986 (courtesy of NOAA, National Marine Laboratory).

Figure 1.7: After the Exxon Valdez oil spill, thousands of birds were killed, and were washed up on Gulf beaches. These sea birds were mainly common murres (Uria aalge) that died offshore (Photograph, courtesy of the Exxon Valdez Trustee Council.) Several hundred thousand sea birds were killed by the spill (see Piatt and Ford, 1996).

Introduction

7

of sockeye salmon abundance during the last 3500 years reveals much about longerterm changes in production, some of several hundred years duration. The Gulf of Alaska has a highly productive food web, culminating in large stocks of crustaceans, fishes, birds, and marine mammals that are exploited for subsistence, commercial and sport fisheries, and form a basis for the growing tourist industry. The harvest of fish and marine mammals can have widespread impacts on the Gulf’s ecosystem. The people of Alaska have historically been an inextricable part of the marine ecosystem. Man, therefore, is both a top predator in the Gulf and, at the same time, uses its waters to transport large quantities of oil and other commodities, including forest products and coal. The results of scientific studies carried out in the northern Gulf of Alaska since the 1989 Exxon Valdez oil spill are refining our views about the response of this ecosystem, its structure, function, and long-term changes in the face of perturbations, both human and natural. Findings from oil spill studies have already advanced our understanding of key ecological relationships, pointed the way to future long-range studies, and contributed to a more informed management of our impacts. In marine research programs and in the development of scientific careers, opportunities for producing synthetic understanding of large ecosystems are relatively rare. Occasionally, after sufficient research has been completed, the time is ripe for synthesis and, possibly, revised concepts about the structure and function of marine ecosystems. The trustees of the publicly held settlement funds from the 1989 oil spill have afforded such an opportunity. So, while the aftermath of the grounding of the Exxon Valdez on Bligh Reef had the dimensions of an ecological and economic tragedy, it also ironically provided the unprecedented means to learn much that will serve the people of Alaska and elsewhere. It is this legacy that is a primary inspiration for our book.

1.1. Why We Have Written This Book Our intent is to synthesize recent information about historical changes in the northern Gulf of Alaska, including those forced by the Exxon Valdez oil spill of March 1989. Our analyses and conclusions are based on a large number of published reports and peer-reviewed manuscripts from a variety of scientific disciplines. We hope that bringing this material together will benefit professionals and non-professionals alike. While our grasp of the causes of change is compromised by the enormous complexity of marine ecosystems as they shift on scales of months, years, decades, and centuries, we believe that “ many small steps make a journey.” There are seemingly competing ideas and theories about the form and function of the marine ecosystem of the northern Gulf of Alaska that we present here. We hope that these ultimately lead to new hypotheses and understandings about the nature of change.

8


1.2. Who Is the Audience for This Book, and What Is Its Scope? The book is intended to serve marine ecologists as well as non-specialists who have a general scientific background and an interest in marine ecology, including, for example, resource managers, graduate students, policy makers, eco-tour leaders and visitors, industry representatives, environmental organizations, and researchers in allied disciplines. We have attempted to present the concepts in the most straightforward and understandable way possible, and to make the material maximally accessible to the wider public – those who do not read the specialized scientific journals where this information appears. At the same time, the authors have not watered down the scientific underpinnings of our knowledge, so it should also be a useful reference for marine scientists seeking a broad synthesis. We have written here about ecological change and what we know of its patterns and causes. This book is not a compilation of everything known about the Gulf of Alaska; there are other references available that are more comprehensive (e.g., Hood and Zimmerman, 1987); nor is it a textbook on oceanography or ecology. We have focused our efforts on ecosystem change as a necessary framework for understanding human impacts because change itself is a primary concern for the people who live next to, and rely on, the Gulf of Alaska.

1.3. Organization The synthesis is presented in five chapters: (1) How does the ecosystem work? In Chapter 2, the ecosystem and its seasonal changes are presented, along with the adaptations for survival in several portal species of fish, seabirds, and mammals. (2) What are the root causes of change? Chapter 3 describes the forces of change – climate, geophysics, species interactions, harvest, disease and contaminants. (3) How has the ecosystem changed in the past? Chapter 4 describes long-term ecosystem changes in physical and biological oceanography, as well as changes in the portal species introduced in Chapter 2 (Pacific herring, pink salmon, pollock, common murres, tufted puffins, black-legged kittiwakes, harbor seals, sea lions, and sea otters). (4) What were the effects of the Exxon Valdez oil spill? The history, fate, and effects of the 1989 spill are described in Chapter 5. (5) What are the reasons for the behavior of the whole system, and what are the emergent properties of ecosystem behavior over the long term? This discussion is presented in Chapter 6.

Introduction

9

References Finney, B. P., I. Gregory-Eaves, M. S. V. Douglas, and J. P. Smol. 2002. Fisheries productivity in the northeastern Pacific Ocean over the past 2,200 years. Nature 416: 729–733. Francis, R. C. and S. R. Hare. 1994. Decadal-scale regime shifts in the large marine ecosystems of the North-east Pacific: a case for historical science. Fish. Oceanogr. 3:4: 279–291. Hanselman, D., J. Heifetz, J. T. Fujioka, and J. N. Ianelli. 2005. Gulf of Alaska Pacific ocean perch. In: Appendix B: Stock Assessment and Fishery Evaluation Report for the Groundfish Resources of the Gulf of Alaska. North Pacific Fishery Management Council, Anchorage, Alaska. pp. 525–578. Hood, D. W. and S. T. Zimmerman (Eds). 1987. The Gulf of Alaska. Physical Environment and Biological Resources. US Minerals Management Service. pp. 655. Piatt, J. F. and P. Anderson.1996. Response of common murres to the Exxon Valdez oil spill and long-term changes in the Gulf of Alaska Ecosystem. In: S. D. Rice, R. B. Spies, D. A.Wolfe and B. A. Wright (Eds.), Proceedings of the Exxon Valdez Oil Spill Symposium, American Fisheries Society Symposium, Bethesda, MD. 18, pp. 720–737. Piatt, J. F. and R. G. Ford. 1996. How many birds were killed by the Exxon Valdez oil spill? In: Proceedings of the Exxon Valdez Oil Spill Symposium S. D. Rice, R. B. Spies, D. A. Wolfe and B. A. Wright (Eds.). American Fisheries Society Symposium. Bethesda, MD, 18, pp. 712–719. Springer, A. M., J. A. Estes, G. B. van Vliet, T. M. Williams, D. F. Doak, E. M. Danner, K. A. Forney, and B. Pfisterer. 2003. Sequential megafaunal collapse in the North Pacific Ocean: An ongoing legacy of industrial whaling? Proc. Natl. Acad. Sci. 100: 12223–12228.


Chapter 2

Ecosystem Structure 2.1. Introduction Robert B. Spies and Alan M. Springer Marine animals thrive in the Gulf of Alaska (GOA) for two basic reasons. First, the ecosystem is productive; a lot of food is available to animals living there. Second, the small number of successful species in the Gulf have made the most of the available energy by organizing themselves, i.e., evolving body forms, physiology, and behaviors that allow them to successfully withstand the rigors of the subarctic ocean, capture food, grow, and reproduce faster than they themselves are eaten, are harvested, or fall victims to contaminants or disease. In this chapter, we will look at the ecosystem from these two perspectives: how the sun’s energy is captured and passed from plants to animals and then from animals to animals; and the strategies that animals have evolved to capture a share of that energy, use it, and flourish in the Gulf of Alaska. The key to understanding long-term change rests on knowing how the ecosystem is “wired,” that is, what is connected to what (the nodes and the connections), but also what it is about each node, i.e., the strategies of successful species that allow them to prosper and dominate. First, to understand why the Gulf of Alaska is so productive, we start with the climate and physical structure of the ocean as it changes through the seasons. In the nine coldest and stormiest months of the year, prevailing winds in the northern Gulf spiral counterclockwise from the central Gulf, pushing surface water onshore and storms into the high ringing coastal ranges. The clouds lose their moisture as they rise over the mountains, dumping large amounts of snow. This precipitation is locked in the cold mountains and coastal plains until spring. As the weather warms, huge amounts of nutrient-poor, but iron-rich, freshwater run off the melting snowpack and glaciers into the nearshore ocean, feeding the Alaska Coastal Current (ACC). The freshwater floats as a layer over the saltier oceanic water near the coast and resists the mixing of deep nutrient-rich water with that at the surface during the summer and early autumn. At this time, the production cycle at most coastal and shelf locations becomes severely constrained by lack of essential nutrients, and the high rates of photosynthesis observed in the early spring decline to low levels. Thus, the northern Gulf of Alaska differs from marine systems off British Columbia and the Washington and Oregon coasts where coastal upwelling occurs over many months. At these more Long-Term Ecological Change in the Northern Gulf of Alaska Robert B. Spies (Editor) © 2007 Elsevier B.V. All rights reserved.

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southerly locations, northerly winds, during the late spring, summer, and early fall, support seasonal coastal upwelling that strongly promotes photosynthesis during the same time the northern shelf in the Gulf is stratified and nutrient-poor in the upper layers. These different physical characteristics tend to predict a relatively unproductive biological system to the north, yet paradoxically the Gulf is very biologically productive. The apparent resolution of this contradiction means that there must be ways other than strong seasonal upwelling for getting deep-water nutrients to the surface to supply plant life at the base of the food web in the growing periods in spring and summer. Finding these ways is a central theme of current oceanographic research on the northern Gulf. As the earth’s orbit first tilts the northern hemisphere towards and then away from the sun, light and temperature increase and then decrease. Subarctic marine organisms have adapted to optimize the production and flow of energy within marine food webs in a widely fluctuating, but generally predictable, seasonal ocean climate. Better understanding the changes in Gulf production between years, decades, and centuries rests on our knowledge of the system viewed from two broad perspectives: (1) how the communities in the Gulf of Alaska are structured and function in response to this annual cycle of physical variability, and (2) how individual organisms use the environment and other organisms during their life cycles to maintain and increase their populations. So, this chapter will explore the changing structure of the ecosystem during a typical year. We will integrate related elements to form a series of stories that begin with climate as the force of seasonal change in the oceanography, move through the annual cycle of biological production, and conclude with survival strategies of key species of fish, birds, and mammals that tie them to their habitats.

2.2. The Physical Environment of the Gulf of Alaska1 Thomas Weingartner The Gulf of Alaska is the semi-enclosed basin comprising the northeast corner of the North Pacific Ocean. Extensive mountain ranges surround all but its southern boundary, which opens onto the North Pacific Ocean and is defined by the northern boundary of the subarctic front. In the 400-km-wide subarctic front, temperatures and salinities decrease rapidly northward (Yuan and Talley, 1996). Although the front shifts, its northern boundary typically lies along 45°N. The western end of the Gulf of Alaska is ill defined, but is 165°W for the purposes of this book. The southern 1

This chapter is based in part on an earlier contribution (Weingartner, 2005), but has been expanded and revised for this synthesis.

Ecosystem Structure 13

boundary of the Gulf thus extends westward from the North American coast for more than 3000 km, so that the Gulf covers an area of about 3.4 × 106 km2 (or 4% of the North Pacific Ocean area). Its small size and open connections to the North Pacific suggest that the Gulf is neither isolated nor insulated from atmospheric and oceanographic processes that occur elsewhere over the North Pacific Ocean. As will be seen, the effects of climatic processes occurring elsewhere in the North Pacific often move into the Gulf by the semi-permanent atmospheric and oceanic circulation pathways. On the other hand, the Gulf cannot be considered as simply a passive recipient of remotely generated climate signals, for these signals are modified regionally as a consequence of the Gulf’s geologic history and its high-latitude setting. Over time, geologic forces have carved a complex coastline and bathymetry that substantially affects the movements and modifications of water and air masses transported into the region from afar. The high-latitude Gulf has relatively cool air and water temperatures, which affects seawater density and the regional hydrologic cycle through partitioning of the atmosphere’s moisture load into snow or rain. Thus, an appreciation of both remote and regional atmospheric and oceanographic processes is required to properly understand the physical processes responsible for the present-day structure of the Gulf of Alaska ecosystem and its response to natural and anthropogenic perturbations. We will first describe the geomorphology of the Gulf of Alaska and then address regional atmospheric conditions. Then the major oceanographic features and the physical processes that influence its marine production will be described. The chapter concludes with a brief description of the atmosphere system of the North Pacific Ocean, emphasizing processes that influence the Gulf of Alaska.

2.2.1. Geomorphology The Gulf of Alaska encompasses several bathymetric provinces, including the abyssal plain, the continental slope, and the continental shelf (including adjoining bays and fjords). The abyssal plain underlies more than 75% of the Gulf’s area and shoals from 4000 m in the southwest Gulf to about 2500 m in the northern Gulf and along the continental rise (Fig. 2.1). Several fracture zones rupture the plain, and numerous seamounts or guyots (with some rising to within 1000 m of the surface) are scattered throughout this domain. Depths shoal rapidly from 2500 m to 250 m, over a distance of 100 km or less, across the steep continental slope, which connects the abyssal plain to the continental shelf. The shelfbreak is between 200 and 300 m deep and delineates the abrupt transition in bottom slope at the juncture of the continental slope and the more gradually sloping continental shelf. The area of the continental shelf is about 3.7 × 105 km2 (Lynde, 1986) or 12.5% of the shelf area of the United States (McRoy and Goering, 1974). The shelf varies in

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Figure 2.1: The bathymetry of the Gulf of Alaska and the topography of the surrounding land (photograph of globe at Geophysical Institute, University of Alaska, courtesy of A. Springer). A = abyssal plain, AR = Aleutian Range, CI = Cook Inlet, CR = Cascade Range, CS = continental slope, KI = Kodiak Island, P = Prince William Sound, QC = Queen Charlotte Islands, SB = shelf break, and SH = Shumagin Islands. width from a minimum of about 5 km near the Queen Charlotte Islands in Southeast Alaska to a maximum of more than 200 km northeast of Kodiak Island. Shelf depths are typically between 150 and 250 m, and bottom depths are often this deep to within even a few kilometers of the coast. The Gulf of Alaska shelf thus differs substantially from most other continental shelves, where bottom depths are generally much shallower than 200 m and which generally shoal smoothly from the shelfbreak to the coast. The Gulf is one of the most tectonically active zones on earth (Jacob, 1987) because the region straddles the convergent Pacific and North American lithospheric plates. The seismic, tectonic, and volcanic activities associated with the intersecting plates are responsible for much of the regional geomorphology, including the system of mountains bounding the coast. These consist of the northward extension of the Cascade Range that arc northward from British Columbia and thence across south–central Alaska and the Aleutian Range along the Alaska Peninsula. Mountain elevations in the eastern and northern Gulf range from 3 to 6 km, whereas those in the Aleutian Range are about 1 km. The mountains are young, rugged, sparsely vegetated, and, because they easily erode, supply abundant sediments to the ocean. The seafloor and coastline are continuously reshaped through faulting, subsidence, landslides, tsunamis, and subsea turbidity flows. As recently as the Great Alaska Earthquake of


1964, portions of the shelf were uplifted by as much as 15 m (Malloy and Merrill, 1972; Plafker, 1972; von Huene et al., 1972). The mountains support numerous glaciers, covering nearly 20% of the Gulf of Alaska watershed (Royer, 1982), and make the region the third largest glacial field on the planet (Meier, 1984). Glacial mass varies through time; in some years, there is a net increase in glacial ice through accumulation of precipitation in the form of ice and/or snow, whereas in other years, glacial mass decreases as the accumulated snow and ice from the past melts and is released into the ocean. Glacial advance and retreat has occurred repeatedly throughout geologic time, but, today, the glacial fields surrounding the Gulf are retreating, and the retreat rate appears to have nearly doubled in the past 25 years (Arendt et al., 2002). Glacial scouring of the underlying bedrock is an important source of fine-grained sediments to the Gulf (Hampton et al., 1987). The major inputs of glacial silt are the Bering and Malaspina glaciers, the Alsek and Copper rivers in the northern Gulf, and the Knik, Matanuska, and Susitna rivers that empty into Cook Inlet in the northwest Gulf (Hampton et al., 1987). The sediment flux is enormous; for example, the Copper, Susitna, and Stikine rivers drain watersheds, which, in aggregate, amount to less than 4% of the area of the Mississippi River basin. Nevertheless, these rivers discharge nearly a third of the sediment load carried by the Mississippi (Wang et al., 1988). The bathymetry of the Gulf of Alaska reflects the diverse tectonic and glacial processes that have operated over the region for millions of years. The fracture zones of the deep basin are aligned along faults and spreading zones within the Pacific plate, and the seamounts are remnants of volcanoes. The shelf bathymetry is particularly complex; the numerous troughs and canyons that cross the shelf are potentially important pathways for water exchange between the shelf and the basin. Extensive channel systems thread through the island archipelago of Southeast Alaska and the Shumagin and Semidi island groups along the Alaskan Peninsula. Subsea embankments and ridges abound as a result of subsidence, uplift, and glacial moraines, and some of these, such as Portlock Bank northwest of Kodiak and Alsek Bank south of Yakutat, are important commercial fishing grounds. Similar geological processes carved the fjords, bays, and headlands along the Gulf’s immensely convoluted coastline. In addition to influencing the ocean circulation, these various geological features provide a diversity of biological habitats. Two of the more prominent embayments are Prince William Sound at the apex of the Gulf and Cook Inlet in the northwest (Fig. 2.2). The sound is about 60 km wide by 90 km long, and has characteristics of a small inland sea (Muench and Heggie, 1978). Numerous islands are scattered throughout this basin, and fjords that are fed glacially thread inland along the perimeter of the rugged coastline. The sound communicates with the shelf through Hinchinbrook Entrance in the east and through several passes in the southwest, with Montague Strait being the most prominent of these. Hinchinbrook Entrance connects the shelf with the sound’s central basin, where depths

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Figure 2.2: The bathymetry of Cook Inlet: Shelikof Strait (left) and Prince William Sound (right). For Cook Inlet: K = Knik Arm, KE = Kennedy Entrance, MC = Middle Cook Inlet; LCI = Lower Cook Inlet; SS = Shelikof Strait; T = Turagain Arm, and U = Upper Cook Inlet. For Prince William Sound: HE = Hitchinbrook Entrance, KP = Knight Island Passage, MS = Montague Strait, and V = Valdez Arm.

exceed 350 m. The northern edge of this basin joins the upper sound through a 300-m-deep trough leading toward Valdez Arm. A second arm curves around to the northwest, where it broadens to form a smaller basin linking the upper sound to Knight Island Passage and the passes along the southwest boundary. Depths exceed 700 m in the small basin at the northern end of Knight Island Passage, and depths within this passage range from 300 to 600 m. In contrast, the shelf immediately south of the sound is shallower and bathymetrically simpler. For example, the shelf is about 120 m deep and flat toward the south of Hinchinbrook Entrance, except in Hinchinbrook Canyon, where depths exceed 200 m. The canyon extends inshore from the shelfbreak to Hinchinbrook Entrance, and, thus, provides a conduit through which continental slope waters can feed the deeper waters of the sound. Cook Inlet extends inland about 275 km from its mouth at the tip of the Kenai Peninsula to its head near Anchorage (Fig. 2.2b). The upper inlet is about 30 km wide and extends northward from the Forelands for 75 km before dividing into Turnagain and Knik arms, each of which protrudes inland an additional 70 km. The upper inlet, including both arms, has depths of 40 m or less and contains extensive tidal flats that are usually exposed at low tide. Lower Cook Inlet is nearly 70 km wide and contains a 100-m-deep channel along its central axis. The mouth of the inlet opens onto the continental shelf at Kennedy Entrance on its eastern edge and Shelikof Strait to the southwest. This strait is a 200-km-long by 50-km-wide rectangular channel, separating Kodiak Island from the Alaska Peninsula, with depths between


150 and 300 m. South of Kodiak Island, the strait’s deep channel veers southeastward and intersects the continental slope west of Chirikof Island. Further southwest, the shelf shallows (100–150 m) and is interrupted by the Shumagin and Semidi island groups. The shelf narrows at its far western end and effectively terminates in Unimak Pass. This narrow (15 km) and shallow (75 m) strait connects the Gulf of Alaska shelf to the Bering Sea (Schumacher et al., 1982; Ladd et al., 2005).

Local Atmospheric Forcing The storm systems that comprise the Aleutian Low Pressure System primarily determine atmospheric conditions in the Gulf of Alaska. Storm frequency and intensity vary seasonally, however, because of the variations in cyclogenesis (low-pressure formation) and under the influence of the Siberian and East Pacific High pressure systems (Wilson and Overland, 1987). Mean monthly maps of sea level pressure (Fig. 2.3) for January, April, July, and October illustrate the seasonality in the strength and

Figure 2.3: Mean monthly sea level pressure fields over the North Pacific Ocean (mb = millibar). The months selected are representative of winter (December–February), spring (March–May), summer (June–August), and fall (September–November).

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position of these pressure systems. From fall through winter, the Aleutian Low deepens, and its center moves westward. Simultaneously, the East Pacific High weakens and migrates southward offshore of California, while pressure builds within the Siberian High as polar air masses accumulate over northeastern Siberia. Occasionally, frigid arctic air spills southward over the Bering Sea and/or the Alaska mainland, spawning storms in the Gulf (Winston, 1955; Businger, 1991). These “polar lows” storms can be particularly dangerous because they are short-lived and difficult to predict, but often include strong winds and sub-freezing air temperatures that can cause vessel icing. From May through August, both the Aleutian Low and Siberian High weaken, while the East Pacific High strengthens and migrates northward to about 40°N. As a consequence of these seasonal changes, storms tend to follow a more southerly path in winter and a more northerly path in summer and fall, with some continuing on into the Chukchi Sea and Arctic Ocean. In all seasons, however, the majority of storms enter the Gulf from the west, where they stall and dissipate because the coastal mountains contain them. Indeed, the frequency with which this occurs led Russian mariners to refer to the region as the “graveyard of lows” (Plakhotnik, 1964). The oceanographic characteristics of the Gulf of Alaska shelf result, to a large extent, from the interaction between the mountains and the storms moving into the Gulf of Alaska. Orographic effects profoundly influence many of the more important physical variables that mold this marine ecosystem, including precipitation rates and types, coastal runoff, wind velocity, and the heat exchange between the atmosphere and ocean.

Precipitation and Runoff Uplift and cooling of moist air by the mountains lead to condensation and heavy precipitation rates around the coast. Measured precipitation rates of 2 to 4 m year−1 are typical throughout the region, although rates in southeast Alaska and Prince William Sound can easily exceed 4 m year−1 (Fig. 2.4). Precipitation rates at higher elevations have not been measured, but are certainly greater because of greater cooling. With the exception of the Alaska Peninsula, coastal precipitation rates are much greater than rainfall rates over the basin and on the shelf. For example, precipitation at Middleton Island, which lies 100 km south of Prince William Sound on the northern shelf, are about 1.5 m year−1 (Danielson et al., in prep., unpublished data), and similar to those over the central Gulf (Baumgartner and Reichel, 1975). Unless stored as snow or incorporated into glaciers, the precipitation swiftly returns to the ocean via the many small streams that drain the steep coastal mountains. Royer (1982) estimates that, on annual average, approximately 24,000 m3 s−1 (750 km3 yr−1) of freshwater enters the Gulf of Alaska shelf between Ketchikan and Seward, with roughly 60% entering Southeast Alaska and the remainder along the south-central coast. The actual amount of discharge onto the shelf is undoubtedly greater because this estimate does


Figure 2.4: Circulation schematic for the Gulf of Alaska, including the basin current structure (North Pacific Current, Alaska Current, and Alaskan Stream) and the Alaska Coastal Current on the continental shelf. The vertical bars indicate the annual precipitation rate compiled from historical precipitation measurements.

not incorporate the discharge from rivers draining interior Alaska (Alsek, Copper, Susitna, etc.), the freshwater discharge along the British Columbian coast, or the altitudinal influence on precipitation rates. Nevertheless, the magnitude and the diffuse or distributed manner in which the discharge enters the Gulf of Alaska is extraordinary when compared to other continental shelf systems, where freshwater enters through one or a few, large “point” sources. In fact, the discharge into the Gulf of Alaska is much greater than the other mid- and high-latitude watersheds, and it exceeds that of the Mississippi River by 40% (Table 2.1). More striking is the disparity in yields; the yield for the Gulf of Alaska is 2–10 times greater than the other watersheds tabulated. Indeed, based solely on this criterion, the Gulf of Alaska coastal watershed has more in common with equatorial watersheds, such as the Amazon River, than it does with mid- and high-latitude drainage basins. Two other sources of freshwater to the continental shelf are the precipitation falling directly onto the shelf and glacial runoff. The former is estimated to be 47 km3 yr−1 (Weingartner et al., 2005), while recent measurements by Arendt et al. (2002) indicate that the glacial meltwater contributes 97 km3 yr−1 to the Gulf of Alaska shelf. In summary, estimates of the freshwater influx to the Gulf of Alaska shelf are uncertain, but probably exceed 1000 km3 yr−1. Moreover, the annual coastal freshwater flux is 50% greater than the total amount of precipitation falling over the Gulf of Alaska but is discharged directly onto the shelf, which contitutes only about 10% of the Gulf’s area.

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Table 2.1: Discharge from some large watersheds.

Watershed Mackenzie Mississippi Yukon Columbia Amazon Gulf of Alaska*

Drainage Area (km2) 1,800,000 3,200,000 850,000 700,000 7,200,000 480,000

Mean Annual Discharge (km3 year–1) 237 536 221 221 5680 750

Yield = Discharge/Area (cm year–1) 13 17 26 32 80 160

*Royer, 1982.

The mean monthly values of freshwater runoff as estimated by Royer (1981a,1982) for the coastal segment between Ketchikan and Seward are given in Fig. 2.5. Discharge varies tremendously through the year with a maximum (~40,000 m3 s−1) in fall and a minimum (~10,000 m3 s−1) in winter when much of the precipitation is stored as snow. Discharge then increases with the onset of spring melt in May and increases more rapidly in August as the rainy season commences. Freshwater discharge to the ocean is important because it affects seawater density. Density is a function of temperature, salinity, and pressure, although the pressure influence is generally small over the depths of interest considered here and will be neglected hereafter. As will become evident later, density gradients (density changes over vertical or horizontal distances) are of fundamental importance in a variety of processes that affect marine ecosystems. Although both temperature and salinity affect density gradients, the salinity effect is generally of greater importance in the Gulf of Alaska because water temperatures vary little throughout the year, except near the surface in summer. Hence, the massive coastal freshwater discharge into the Gulf of Alaska modifies ocean salinities and, in so doing, profoundly affects vertical and horizontal density gradients. Vertical density gradients, or ocean stratification, affect biological production and mixing. Horizontal gradients are indicative of horizontal pressure gradients that force horizontal currents. (We will refer to such currents in this book as density gradient flows.) While the ocean’s density structure is strongly dependent on the discharge, winds and temperature changes induced by heat exchanges between the atmosphere and ocean also affect the density structure. Winds modify density gradients through vertical stirring (mixing) and through advection of waters from elsewhere.


Figure 2.5: Mean monthly coastal freshwater discharge (after Royer, 1981a, 1982) and mean monthly wind speed from Middleton Island on the northern Gulf of Alaska shelf. Negative wind speed values imply that the mean winds are westward.

Winds The winds associated with low-pressure systems over the Gulf of Alaska blow counterclockwise (with the coast to the right of the direction of the wind), and frictional coupling between wind and water (stress) forces the ocean circulation. Wind stress is proportional to the square of the wind speed and forces a surface Ekman transport (typically confined to the upper 10 to 30 m of the water column), which is directed 90° to the right of the wind in the Northern Hemisphere. On the continental shelf, the surface Ekman transport induces a sea level rise (relative to offshore) along the coast and sinking (downwelling) of surface waters within a narrow coastal band. The cross-shore sea level slope is small (amounting to a difference of only a few centimeters over 100 km), but it results in a cross-shore pressure gradient that, in conjunction with the earth’s rotation, propels a steady along-shelf current in the downwind direction throughout the water column. Because most of the surface Ekman transport sinks along the coast, there must be a compensatory offshore flow at depth (otherwise the sea level would continue to rise along the coast for as long as the wind blew). Under relatively steady wind forcing, the offshore flow is contained within a bottom Ekman layer established by frictional stress between the along-shelf flow and the seabed. An alongshore wind blowing with the coast to its left is an upwelling-favorable wind. Under such conditions, coastal sea level decreases, and the alongshore flow and surface and bottom Ekman transports are opposite in direction from the

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downwelling case. Such winds also induce an uplift of deeper waters to the surface within a narrow coastal band; a process referred to as upwelling. As will be seen later, vertical and horizontal differences in water density alter this circulation structure in important ways, although the basic features outlined earlier are nonetheless retained. As an example of the annual cycle in wind forcing over the northern Gulf of Alaska, Fig. 2.5 shows the mean monthly along-shelf wind velocities measured at Middleton Island. The monthly averages, based on long-term measurements at Middleton, show that downwelling winds prevail throughout the year and are maximum in winter and minimum in summer. Spatial and temporal variations in wind stress complicate the circulation response outlined earlier, however, and the Middleton Island winds are not necessarily representative of conditions around the shelf. Unfortunately a detailed understanding of the temporal and spatial wind field over the Gulf of Alaska has yet to be achieved, because, with the exception of Middleton Island, direct, long-term wind measurements on the continental shelf and basin are scarce. Although there are numerous coastal communities in which winds have been measured for many years, these measurements are influenced by local orographic effects that make extrapolation of coastal measurements onto the shelf problematic. Our present understanding of spatial and temporal variations of the shelf wind field is instead derived from limited-duration measurements obtained from moored buoys and/or inferred from wind estimates generated by weather forecast models. In a general sense, these indicate that temporal variations are correlated over broad areas of the Gulf (especially in fall and winter, when winds are strongest) because storms entering the Gulf are comparable to the size of the basin itself (Livingstone and Royer, 1980) (Fig. 2.6). This result does not imply, however, that the wind field is spatially uniform around the shelf. Indeed, it appears that whereas, on monthly average, downwelling-favorable winds blow over the northern and eastern Gulf of Alaska year-round, the maximum in wind stress occurs in the northeast Gulf. Fall and winter wind directions are more variable in the western Gulf because storms entering the Gulf follow two predominant pathways: a west – east track across the Alaska Peninsula from the Bering Sea or a southwest – northeast track across the Gulf of Alaska. As a result, winter wind stress over the shelf between Kodiak and Unimak Pass appears to be upwelling favorable on average (Stabeno et al., 2004). Winds weaken throughout the Gulf in summer and, while downwelling winds prevail over the northern Gulf, upwelling-favorable winds occur along the British Columbian shelf and southwest Gulf of Alaska. While the mean monthly wind estimates provide a useful description of the annual cycle, these estimates belie the episodic nature of wind forcing, which is tied to storm systems and weather fronts propagating across the Gulf every 3 to 10 days. Particularly in summer, there are prolonged periods of relative calm or weak, upwelling-favorable winds that are interrupted by relatively strong downwelling winds associated with only a few storms of several days duration in each month.


Figure 2.6: Low-pressure system in the Gulf of Alaska. The MODIS instrument aboard NASA’s Aqua satellite captured this true-color image of a large low-pressure system spinning in the Gulf of Alaska on August 17, 2004. The spinning cloud systems of these lows express strong, persistent, low-pressure systems that form because of the persistent flow of semi-permanent pressure systems north (the polar easterlies) and south (the subtropical high) of the area. Image courtesy Jesse Allen: http://earthobservatory.nasa.gov/NaturalHazards/Arch.

Important wind stress variations also occur at smaller (mesoscale) spatial scales (30–100 km). These are especially prominent in winter and are closely tied to orographic effects and the channeling influence of straits and the mouths of bays and fjords. Mesoscale winds are enhanced in winter when the coastal mountains inhibit mixing between the cold, dry air of Alaska’s interior and the warmer, moister marine air. Mountain passes provide a conduit through which continental air pours across the shelf and collides with marine air (Macklin et al., 1988). These cold-air outbreaks produce offshore-directed gap winds at the mouths of coastal embayments, including Hinchinbrook Entrance, the Copper River delta (Macklin et al., 1988), Shelikof Strait (Lackmann and Overland, 1989; Macklin et al., 1984), lower Cook

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Inlet, and the Barren Islands (Macklin et al., 1990). Under similar winter conditions, the coastal mountains can also enhance the alongshore wind component under downwelling-favorable situations. In this case, the shelf wind field includes a narrow (50–100 km wide) coastal barrier jet (Parish, 1982; Overland and Bond, 1995) having wind speeds several times greater than those further offshore. Mesoscale wind fields affect the coastal ocean in two significant ways. First, the barrier jets and the largergap wind events (such as those flowing through Shelikof Strait and/or offshore of Cook Inlet) can substantially affect the cross-shelf and along-shelf current structure. (The smaller-scale-gap winds at the mouth of bays are unlikely to have a significant influence on the shelf circulation, although they could be locally important in influencing exchange between the shelf and the bay or fjord.) Second, winter mesoscale winds are likely important in cooling the coastal ocean because the rate at which heat is lost from the sea is a function of both wind speed and the air–sea temperature difference. Although Overland and Heister (1980) suggest that these wind events are common in winter, the frequency of these phenomena and their importance in the coastal ocean has not been quantified. It also seems probable that the frequency and intensity of mesoscale wind phenomena will vary from year to year in accordance with atmospheric conditions in the Alaskan interior and the Gulf.

Heat Fluxes Changes in upper-ocean temperatures over both the basin and the continental shelf are largely controlled by the exchange of heat between the atmosphere and ocean. This exchange depends on a number of variables, including the air and sea temperatures and their difference, atmospheric humidity, wind speed and direction, and the sky or cloud cover. Sky cover affects the amount of radiation available to heat the upper ocean and sustain marine photosynthesis. Clouds form as moist air rises, expands, and cools, either during cyclogenesis or over the coastal mountain ranges. Sky cover over the Gulf of Alaska is extensive and varies little seasonally. Indeed, clouds obscure 60% or more of the sky more than half of the time, while a sky cover of 25% or less occurs only about 15% of the time (Brower et al., 1988). Because of its high-latitude location, the influx of solar radiation into the Gulf varies from a maximum of about 225 W m−2 at the summer solstice to about 10 W m−2 at the winter solstice (Fig. 2.7). (Over a column of water 30 m deep and 1 m2 in surface area, solar radiative heating at the summer solstice would lead to a 1.6°C rise in temperature in 10 days, whereas the winter solar radiative flux would increase temperatures by 0.08°C over the same period.) Interannual variations in skycover, especially in summer, can affect upper-ocean temperatures. As an extreme but not unrealistic example, a 25% reduction in summer cloud cover could increase temperatures in the upper 30 m of the water column (where most of the solar radiation is absorbed)


Figure 2.7: Mean monthly heat fluxes for short-wave, long-wave, latent, sensible, and the net or the sum of all terms for the Gulf of Alaska.

by 1–2°C. Although this temperature increase appears slight, it could be important to marine invertebrates whose metabolic rates are temperature dependent. Heat is lost from the ocean to the atmosphere by sensible and latent heat exchange and long-wave radiation. Each of these processes causes cooling in the Gulf of Alaska throughout the year (Fig. 2.7). The heat transfer rate for latent and sensible heat exchanges increases as both the air–sea temperature contrast and the wind speed increases. Long-wave, radiative heat loss occurs because the ocean surface emits radiation to the atmosphere at a rate proportional to the surface temperature, although the effectiveness of this cooling is reduced by clouds. As with much of the earth’s surface, the ocean radiates at infrared wavelengths that, on absorption by water vapor, carbon dioxide, and other gases, lead to atmospheric warming, e.g., the “greenhouse gas” effect. While long-wave heat loss is relatively invariant from month to month, the sensible and latent heat fluxes vary substantially throughout the year and are greatest in winter (when air–sea temperature differences and wind speeds are greatest) and least in summer. The net heat flux (based on the sum of all the heat fluxes) results in a distinct annual heating and cooling cycle that coincides with the solar cycle. Thus, the ocean gains heat between the spring and fall equinoxes and loses heat to the atmosphere over the rest of the year (Fig. 2.7). Averaged over the year, there is a net heat loss of 13 W m−2 over the northern Gulf of Alaska shelf to the atmosphere, as estimated from Middleton Island data. Most likely, the nearshore winter air temperatures are cooler

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and wind speeds higher than those measured at Middleton, so that inner-shelf winter heat losses are probably appreciably greater. Consequently, inner-shelf waters (including Prince William Sound, Cook Inlet, and Shelikof Strait) are generally colder than outer-shelf waters in winter. Nevertheless, the annual imbalance between summer heat gain and winter heat loss based on Middleton Island data suggest an annual decrease in shelf water temperatures of at least 0.8°C over the upper 100 m. This implies that on long-term average, the ocean circulation must advect into the Gulf the heat required to balance that lost annually to the atmosphere.

2.2.2. Physical Oceanography The Seasonal Cycle of Water Properties The annual cycles of wind, freshwater inputs, and air–sea heat flux are reflected in seawater temperatures and salinities on the shelf. These changes can be seen in the 2.5-year record of hourly temperature and salinity values (Figs. 2.8 and 2.9) collected at the hydrographic station GAK 1, which lies at the mouth of Resurrection Bay offshore of Seward. The plots accompanying the mean monthly values show the amplitude and phase for the annual cycle as a function of depth based on a least squares fit to the annual cycle. (The leftmost plot shows the mean annual temperature and the range. The middle plot shows the phase or the time of the year when the maximum temperature or salinity occurs. The rightmost plot shows the variance accounted for by the annual period.) The annual period accounts for 90% (at the surface) to 50% (at the bottom) of the variance for both temperature and salinity (Weingartner et al., 2005). Mean sea-surface temperatures range from a minimum of 3.5°C in March to 14°C in August. The range of temperatures decreases with depth, however, and is less than 1°C annually at depths greater than 150 m. Although the lowest sea surface temperatures coincide with the end of the cooling season, the highest surface temperatures occur in August, about one month before the end of the heating season in late September (Fig. 2.7). Surface temperatures begin to decline in September because increasing wind-mixing and downwelling redistribute heat over the upper portion of the water column. These processes lead to the development of a subsurface temperature maximum in October. From fall through early winter, temperatures within the subsurface maximum temperature decrease, and the depth of this layer increases such that bottom water temperatures attain their annual maximum in January. Bottom water temperatures exceed those at shallower depths from January through April. Temperatures decline throughout the water column through April because of deep winter mixing and vigorous surface cooling. By late April or early May, near-surface temperatures begin to increase as day length increases, so that the water column becomes nearly isothermal in early May. Through midsummer, temperatures increase rapidly at the surface and more slowly at mid-depth, while temperature remains nearly constant at depths greater than 100 m.


Figure 2.8: Water temperatures (December 1999–March 2002) at selected depths (upper panel) at the hydrographic station GAK 1 near Seward, Alaska. The lower panel shows the results of a least squares fit to the annual harmonic for each calendar year (2000 and 2001) at each measured depth. The leftmost panel shows the mean (solid circle) and amplitude (horizontal bars), the middle panel, the phase (time of the maximum temperature at the specified depth), and the rightmost panel, the fraction of the total variance explained by the annual harmonic (from Weingartner et al., 2005).

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Figure 2.9: Water salinities (December 1999–March 2002) at selected depths (upper panel) at hydrographic station GAK 1 near Seward, Alaska. The lower panel as in Fig. 2.8 (from Weingartner et al., 2005).


Surface water is saltiest in April and freshest in August. The range of surface water salinity is from 31 to about 25 PSU (practical salinity unit). Although surface salinity begins to increase in September, salinities in the upper 50 m decrease through early November because wind mixing and downwelling redistribute salt and freshwater over the upper layers. Consequently, the upper-layer salinities vary nearly in-phase with the annual discharge cycle. At depths greater than 100 m, however, minimum salinities occur in late winter following maximum downwelling, and maximum salinities occur in summer when downwelling is weak. Thus, the annual cycle in salinity over the lower half of the water is out-of-phase with the surface salinity cycle but in-phase with the annual cycle in along-shelf winds. The summer salinity increase below 150 m heralds the arrival of salty, nutrient-rich water from offshore that spreads inshore from the continental slope (Royer, 1975; Xiong and Royer, 1984; Weingartner et al., 2005). The near-bottom winter salinity minimum results from strong fall and winter downwelling, which gradually flushes deeper, more-saline waters offshore and/or mixes them with fresher surface waters. The annual temperature and salinity cycles described in the preceding text are characteristic of the Gulf of Alaska shelf in general, although there are important spatial differences. For example, at distances of more than 50 km from the coast, the annual range in salinity is generally about 1 PSU or less at all depths. Surface salinities in glacial fjords and bays (e.g., Cook Inlet) or near the mouths of rivers entering the shelf can be substantially fresher than elsewhere along the coast. In general, winter surface water temperatures are cooler along the coast than further offshore for the reasons discussed earlier. There are also alongshore gradients in these variables; for example, winter surface water temperatures decreasing along the coast and continental slope between Southeast Alaska and the Alaska Peninsula. Seasonal changes in temperature and salinity in the water column affect the density stratification, which in turn affects the extent of vertical mixing. A strongly stratified water column requires more energy to mix than a weakly stratified one. Seasonal changes in stratification affect photosynthesis by phytoplankton, which require both light and nutrients. In general, plant nutrient concentrations increase with depth, while the photosynthetically important wavelengths of light are rapidly attenuated with depth. When stratification is weak, deep mixing removes phytoplankton from the euphotic zone, the depth layer in which there is sufficient light for plant growth. On the other hand, this mixing supplies the euphotic zone with nutrients. The water column is weakly stratified from January through March, so deep mixing enriches the surface layers with nutrients. The decrease in wind speeds and the increase in runoff and solar heating in April and/or May trigger stratification. When the stratification is strong enough to prevent mixing to depths greater than the euphotic zone depth, a spring bloom of phytoplankton occurs (see Section 2.3 for a more detailed explanation and figures). Although stratification is building in early spring, storms can interrupt the process by inducing deep mixing. Nevertheless, by late May, the

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shelf is generally stratified everywhere, with stratification strengthening through summer. Consequently, from late spring through early September, the shelf and basin consists of a shallow, well-stirred, surface-mixed layer about 20 m thick, over which salinity and temperature are nearly constant. The summer mixed layer is separated from deeper layers by a thin, vertical layer across which density changes very rapidly (the pycnocline). The summer pycnocline is typically between 15 and 40 m in depth and coincides with both a thermocline and a halocline. While nutrient concentrations are nearly depleted in the summer mixed layer due to phytoplankton consumption in spring, high concentrations are found just beneath the pycnocline. Stirring of the water column, by winds, entrainment of deeper water into a spreading plume of low-density water, or tidal mixing processes, can inject nutrient-rich water into the euphotic zone and restimulate plankton production. As winds intensify in fall, the stratification erodes due to both stronger vertical mixing and, at least near the coast, increased downwelling. There are important spatial variations in the onset of springtime stratification and in the structure of the seasonal pycnocline on the Gulf of Alaska shelf. The inner-shelf, bays, fjords, and Prince William Sound stratify first because coastal freshwater runoff is confined initially to nearshore regions and only gradually spreads offshore, if at all, through ocean advection and horizontal mixing processes. Solar heating provides additional surface buoyancy by warming the upper ocean layers over both the shelf and basin. However, thermal stratification often remains weak until May, so that the inner shelf stratifies earlier than the outer shelf and the northern Gulf of Alaska basin. While the springtime onset of stratification on the inner shelf depends primarily on freshwater runoff, surface heating primarily controls seasonal variations in stratification over the outer shelf, slope, and basin. Spatial or temporal variations in the timing, strength, or frequency of stratifying processes (warming, freshwater runoff) and destratifying processes (winds, tidal mixing) consequently affect phytoplankton production and, thus, the amount of food energy available for higher trophic levels. The annual cycle of nutrients is intimately linked to seasonal changes in water column stratification, temperature, and salinity. Childers et al. (2005) described the annual cycle of nitrate, silicate, phosphate, and ammonium across the shelf. These changes are similar to those at station GAK 1, as shown in Fig. 2.10. Uniform nutrient concentrations occur throughout the water column in March prior to the onset of the spring bloom. Except for ammonium, these are rapidly drawn down by phytoplankton consumption in the upper 50 m by late May and remain low through October. Upper-level ammonium concentrations increase in spring by zooplankton excretion. Deep nutrient concentrations tend to follow salinity changes, with maximum concentrations occurring in summer.

Circulation over the Gulf of Alaska Shelf and Slope In a broad sense, the major circulation features of the Gulf of Alaska circulation (shown schematically in Fig. 2.4) consists of a system of counterclockwise flows on


Figure 2.10: Seasonal cycle of the macronutrients: silicate (Si[OH]4, top); nitrate (NO3, middle); and phosphate (PO4, bottom) at selected depths at station GAK 1, at the mouth of Resurrection Bay, near Seward.

the shelf, slope, and central basin. There are, however, substantial differences in the structure of the currents and the physics controlling these various flows, and, for this reason, the circulation in each region is described separately in the following subsections. Flow on the inner shelf consists of the Alaska Coastal Current (ACC), while it includes the Alaska Current along the continental slope (in the eastern and

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northeastern Gulf) and its transformation into the Alaskan Stream (in the northwestern Gulf). Both these current systems are swift (compared to flow in the central basin) and extend over a broad alongshore domain. By the nature of their geographical extent, these currents are likely important in conveying climate perturbations around the Gulf and, ultimately, into the Bering Sea. Although the circulation within the central basin is much feebler than that of shelf or slope, it involves a large mass of water and, hence, climate signals introduced into the basin interior may remain for a relatively long time. Each of these current systems and their advected water masses respond on seasonal and longer time scales to winds, coastal freshwater discharge, heat exchange with the atmosphere, and the advection of water masses formed elsewhere in the North Pacific Ocean. The water masses and circulation in the Gulf of Alaska shelf define three domains: the Alaska Coastal Current domain on the inner shelf, a mid-shelf region that extends from the offshore edge of the ACC to near the outer shelf, and an outer shelf region, which includes the shelfbreak and continental slope. These regions are separated from one another by frontal systems that act as semipermeable boundaries limiting exchanges of mass, material, and momentum. Frontal strength varies in accordance with the magnitude of horizontal density gradients, which, in the Gulf of Alaska, is primarily a consequence of salinity gradients. Fronts also support a three-dimensional flow field. However, the strongest component of flow is parallel to the front and oriented such that low-density water is located to the right of the flow direction in the Northern Hemisphere. As a preface to subsequent discussions, we present schematics of the circulation and frontal structure on the Gulf of Alaska shelf for fall through spring and late spring through early fall (Fig. 2.11). From fall through spring, relatively strong alongshelf winds impel an onshore Ekman transport of surface waters and downwelling near the coast. This process traps low-salinity waters (freshened by coastal runoff) on the inner shelf, forming a deep front within 35 km of the coast. The main axis of the ACC coincides with the front, and the region inshore of and including the front constitutes the ACC domain. A second front forms at the shelfbreak and separates the mid-shelf domain from the outer shelf, which includes the Alaska Current/Stream flowing along the continental slope. The shelfbreak front, which may or may not have a surface expression, is anchored to the shelfbreak and inclined in the offshore direction. Weak upwelling occurs on the inshore side and at the base of the front (Gawarkiewicz and Chapman, 1992; Pickart, 2000). In summer, when runoff is increasing but winds are weak, the ACC front is shallow and may spread 50 km or more offshore. With the reduction in downwelling (or an increase in upwelling), the alongshore flow over the shelf and slope weakens, causing the foot of the shelfbreak front to shoal and reattach at a shallower isobath inshore (Chapman, 2000; Weingartner et al., 2005). The cross-shore flow fields associated with surface and bottom Ekman transports also change seasonally. In winter, the surface onshore Ekman transport is compensated


Figure 2.11: Schematic of the circulation in fall, winter, and spring (left) and summer (right) over the shelf and slope of the northern Gulf of Alaska. The view is looking westward with the coast (and north) to the right. Cross-shelf flow and vertical flows are indicated by arrows. Flows over the middle shelf are variable, although weakly westward on average.

for by an offshore flow along the bottom and/or within the interior. In summer, the onshore Ekman transport is much weaker, but there is also an onshore flow at the bottom associated with the shoreward movement of the base of the shelfbreak front. A compensatory return flow presumably occurs in the interior, although how this is accomplished on the Gulf of Alaska shelf is not understood. Although these are the major seasonal differences in frontal structure over the Gulf of Alaska shelf, intraseasonal frontal variations can be substantial due to fluctuations in winds, alongshore variations in bathymetry, meanders in the ACC and slope flows, and the passage of large eddies along the continental slope. The magnitude of current speeds in the along-shelf, cross-shelf, and vertical directions are quite different from one another. Cross-shelf currents associated with Ekman transports are only a few cm s−1 and smaller (by a factor of 10–100) than the alongshore shelf currents. Vertical current speeds are even smaller, being a few meters or even less per day, except in the strong downwelling regions near the coast during intense storms when vertical speeds can be as large as 10–20 m/day.

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The Alaska Coastal Current (ACC) The ACC is the most prominent aspect of the shelf circulation (Royer, 1981b; Johnson et al., 1988; Stabeno et al., 1995). Although slender, having a typical width of 35 km, it extends for over 2500 km from its probable origin on the British Columbian shelf, although in some months or years it might originate as far south as the Columbia River (Thomson et al., 1989; Hickey et al., 1991; Royer, 1998) to Unimak Pass in the western Gulf, where it leaves the Gulf of Alaska shelf and enters the Bering Sea (Schumacher et al., 1982, Ladd et al., 2005). The properties of the ACC in Southeast Alaska are poorly known, but it seems likely that much of it flows through the complex of channels threading through the islands. The current then continues westward along south-central Alaska until it reaches Prince William Sound. Much (if not most) of the ACC flow enters the sound through Hinchinbrook Entrance and exits through Montague Strait (Niebauer et al., 1994). The remainder of the current continues across the mouth of Hinchinbrook Entrance, southwestward along Montague Island, and thence westward after rounding the southern tip of this island. West of Montague Island, this branch of the ACC and the outflow from Montague Strait merge to continue westward along the south coast of the Kenai Peninsula. The ACC apparently splits northeast of Kodiak Island (Stabeno et al., 1995), with some of the current flowing southward along the shelf east of Kodiak Island. Most of the current curves around the mouth of Cook Inlet, however, before continuing southward through Shelikof Strait (Muench et al., 1978). As it arcs across the mouth of the inlet, bottom topography induces upwelling and locally strong tides mix salty, nutrient-rich water to the surface. Some of this upwelled water flows northward, supplying nutrients and salt (and possibly heat in winter), along the eastern shore of Cook Inlet. The inflow is gradually mixed by tides with freshwater from the rivers that enter along the sides of the inlet to form a dilute southward flow along the west side of Cook Inlet. This outflow then rejoins the ACC at the head of Shelikof Strait, with the mixture continuing flowing southward through this channel. At the lower end of Shelikof Strait, the current continues along the Aleutian Peninsula and through the Shumagin/Semidi island group en route to Unimak Pass. The seasonal transitions implied by the schematics (Fig. 2.11) are evident in the cross-shelf distribution of temperature and salinity properties shown in Fig. 2.12 for April, August, and October. (The data were collected on a cross-shelf transect on the northern Gulf of Alaska shelf offshore of Seward.) In winter (February through April), both the vertical and cross-shelf gradients of salinity and temperature are weak. The ACC front lies within about 10 km of the coast and extends from the surface to the bottom. In summer, the vertical stratification is large but cross-shelf salinity (and density) gradients are weak. At this time, the ACC front extends 50 km offshore and is usually less than 40 m deep. Vertical stratification weakens in fall, although the cross-shelf salinity gradients and the ACC front are stronger than at other times of the year. As along-shelf winds and coastal downwelling increase in


Figure 2.12: Cross-shore distributions of temperature (left) and salinity (right) in the northern Gulf of Alaska, contoured as a function of horizontal distance offshore (horizontal axis) and pressure (vertical axis). Pressure units are approximately equal to the depth in meters. The view is eastward with the coast (and north) on the left. Waters of lowest salinity are usually within 35 km of the coast and represent the Alaska Coastal Current.

fall, the front moves shoreward to within 30 km of the coast and steepens so that the base of the front intersects the bottom between the 50- and 100-m isobaths. The front deepens and steepens through January, but temperature and salinity gradients gradually weaken in response to wind mixing, surface cooling, and the reduction in runoff. A further notable feature in these figures is that cross-shelf temperature differences at all depths are generally small and rarely exceed 1–2°C. In contrast, vertical changes in temperature can be quite large, especially in summer, when a strong thermocline centered at about 25 m depth forms across the shelf. Temperatures decrease from 14–8°C across the summer thermocline. The structural changes in the ACC throughout the year are also accompanied by changes in the velocity distribution and transport. The large cross-shore density gradients

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within the ACC front are associated with currents having substantial current shear. Speeds within the ACC front often exceed 30 cm s−1, but can approach 200 cm s−1 (Johnson et al., 1988) in fall, while speeds elsewhere in the ACC domain are generally 5–20 cm s−1. The shears in summer are primarily confined to the upper 50 m, with weaker (and possibly even reversed or eastward) flow throughout the rest of the water column. Along-shelf velocities increase in fall when substantial shears develop over the upper 100 m of the water column. In winter, when the current’s dynamics are primarily winddriven, vertical current shears are weak and speeds are more uniform throughout the ACC domain. Although maximum speeds are observed in the fall, it appears that maximum transport is in winter (Schumacher et al., 1989; Stabeno et al., 1995), while minimum transport occurs in summer. Stabeno et al. (1995) estimate that, on annual average, the ACC transports at least 800,000 m3 s−1, although transport variations of several days to a week can be large and exceed 3,000,000 m3 s−1. The complex vertical and cross-shore circulation cells within fronts often result in enhanced biological production and the trapping of materials within the front (Garrett and Loder, 1981; Yankovsky and Chapman, 1997; Chapman and Lentz, 1994; Chapman, 2000, Williams, 2003). This appears to be true for the ACC front as well. For example, surface drifters released seaward of the ACC front move onshore (in accordance with Ekman dynamics) and then westward upon reaching the front (Royer et al., 1979). Conversely, the surface layer spreads seaward on the inshore side of the front with the rate of offshore flow increasing with the discharge (Johnson et al., 1988; Williams, 2003). Together, these results suggest cross-frontal flow convergence arising from differing dynamics on either side of the ACC front. Freshwater runoff forces offshore surface flow on the inshore side of the front (with this influence varying throughout the year in accordance with the runoff), whereas wind forcing dominates offshore of the front. One consequence of this flow structure is that crossfrontal exchange of water and dissolved and suspended materials (including plankton) are inhibited (Csanady, 1984). A second consequence is that plankton might accumulate along the frontal boundary, possibly attracting foraging fish, seabirds, and marine mammals. The summer and fall current shears embedded in fronts might also affect predator–prey interactions. Phytoplankton, juvenile salmon, and forage fishes feed in the upper 20 m of the water column at this time of the year (Boldt, 2001; Coyle and Pinchuk, 2003) and drift with the current. However, the zooplankton, which feed on the phytoplankton and on which the fish prey, migrate diurnally over at least the upper 100 m. Hence, diurnally migrating zooplankton are unlikely to encounter the same phytoplankton patches and fish schools over a day because of the highly sheared flow.

Mid-shelf domain The mid-shelf domain is bracketed inshore by the ACC front and offshore by the shelfbreak front. The position of these fronts varies on seasonal and shorter time


scales, so that the dimensions of the mid-shelf domain also vary. The width of the mid-shelf domain will also vary in accordance with the shelf width. In the northwest Gulf of Alaska, the mid-shelf domain is about 100 km wide, but on the narrow shelf of the eastern Gulf, the mid-shelf domain might be very narrow and possibly nonexistent. Stratification over the middle of the shelf is always weaker than within the ACC, except in late winter, when both regions are weakly stratified. As with the ACC domain, salinity controls mid-shelf stratification from fall through late winter, but, in contrast to the ACC, temperature controls stratification in spring and summer. In general, crossshelf temperature and salinity gradients are relatively weak throughout the year, so that the density gradient component of flow over the mid-shelf is weaker than inshore, although still westward on average (Niebauer et al., 1981; Hermann et al., 2002). Flow over the mid-shelf domain is poorly understood, however, having received comparatively little study. Nevertheless, existing observations and models indicate that the mid-shelf domain contains energetic current variations, involving both current reversals and strong cross-shelf flows. The most likely sources of this variability are mesoscale (10–50 km) flow features and, less frequently, large (~150 km diameter) eddies that impinge on the continental slope. Mesocale variability originates in several ways, including, (1) separation of the ACC from coastal headlands with subsequent eddy formation (Ahlnäes et al., 1987; Cenedese, 2002), (2) eddy shedding or meandering by the ACC (Mysak et al., 1981; Bograd et al., 1994), (3) interactions of the shelf flow with variations in bottom bathymetry (Lagerloef, 1983), and (4) meandering of the shelfbreak front along the continental slope (Niebauer et al., 1981) induced by fluctuations in the Alaska Current/Stream. It is, however, extremely difficult to quantify mesoscale variability because it varies both spatially (depending on coastal landforms, bottom topography, and ambient currents) and temporally (depending on the winds and the shelf and slope density distribution). The transient but energetic mid-shelf currents are undoubtedly important to the shelf ecosystem by facilitating cross-shelf exchanges of organisms and dissolved and suspended materials. The mid-shelf thus serves as an important link between the inner shelf and the continental slope. Some of the eddies and current meanders include vigorous vertical motions capable of altering mixed layer depths and advecting nutrients into the euphotic zone. Moreover, until eddies dissipate, they do not readily exchange mass and materials with ambient waters and so may serve as important nurseries for some organisms. For example, in Shelikof Strait, the abundance of larval pollock is generally greater in eddies than outside them (Incze et al., 1989; Vastano et al., 1992; Schumacher and Kendall, 1991; Schumacher et al., 1993; Bograd et al., 1994). Moreover, juvenile pollock collected within eddies are in significantly better condition than those collected outside (Canino et al., 1991), consistent with these features enhancing production of both phytoplankton and zooplankton. Although the linkages between the inner and outer shelf are incompletely understood, the annual inshore incursion of the foot of the shelfbreak front is a robust

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feature of the summer season, as discussed with respect to the schematic in Fig. 2.11. It is clearly evident in the August panel of Fig. 2.12 and delineated by the 32.8 salinity isopleth. The influx of high-salinity water probably exerts an important dynamical influence on the shelf circulation by modifying the stratification above the shelf bottom boundary layer (Gawarkiewicz and Chapman, 1992; Chapman, 2000; Pickart, 2000). The deep onshore flow might also serve as a pathway for deep oceanic zooplankton. Some of these organisms migrate diurnally over the full depth of the water column and only visit the surface to feed for a short period at night. Hence, the bottom flow that transports the high-salinity water shoreward might also result in a net shoreward flux of zooplankton in summer. The summertime inflow of saline water onto the inner shelf also allows the slope and basin interior to communicate directly with the nearshore, for (as discussed in the following text) this water is drawn from within the permanent halocline of the Gulf of Alaska. Finally, the deep summer inflow is a potentially important conduit for nutrients from offshore to onshore, although it is not the only means by which nutrient-rich offshore water supply the shelf. Other mechanisms include flow up canyons that intersect the shelfbreak (Klinck, 1996; Allen, 1996, 2000; Hickey, 1997), topographically-induced upwelling (Freeland and Denman, 1982), and shelfbreak eddies and flow meanders (Bower, 1991). Reed et al. (1987) and Schumacher et al. (1989) suggest that as the ACC flows southward through Shelikof Strait, it entrains deep salty and nutrient-rich water from the canyon depths into the surface layer. The deep waters within the canyon are then replenished by a deep inflow into the mouth of the canyon at the continental slope south of Kodiak.

Prince William Sound The circulation and water mass structure of Prince William Sound is intimately linked to the Gulf of Alaska shelf primarily through the ACC, although there are also conspicuous linkages to offshore waters, as discussed earlier. The basic circulation pattern within the sound appears to be as follows. Waters enter the sound through Hinchinbrook Entrance and then flow counterclockwise around the central basin. Some of this flow passes immediately into the western sound and exits through Montague Strait (and also perhaps along the western side of Hinchinbrook Entrance), and some of it continues into the northern sound (Schmidt, 1977; Royer et al., 1990; Niebauer et al., 1994; Gay and Vaughan, 2001). Northern sound waters flow southward through Knight Island Passage and reenter the shelf through the passes in the western sound, including Montague Strait. This circulation pattern varies seasonally in accordance with the seasonal cycle of winds and runoff and appears to be strongest in late fall and winter and weakest in summer. Indeed, the counterclockwise circulation pattern might even reverse occasionally, if not consistently through summer, with surface waters leaving through Hinchinbrook Entrance and entering through Montague Strait (Vaughan et al., 2001). Although there are no precise estimates of mass and material exchanges between


the sound and shelf, Niebauer et al. (1994) suggest that nearly half of the volume of the sound above 100 m depth is exchanged between May and September, whereas this same layer is renewed at least twice between October and April. There is also deep exchange of waters between the shelf and the sound. This occurs primarily through Hinchinbrook Entrance and also varies seasonally. A deep inflow of salty, nutrient-rich waters occurs in summer, coincident with the inshore migration of this slope water mass elsewhere across the shelf bottom. This flow is channeled northward across the shelf through Hinchinbrook Canyon and is probably an important contribution to the annual nutrient budget of the sound. In winter, fresher waters (formed by deep winter mixing) leave the sound in the lower layer (Niebauer et al., 1994; Vaughan et al., 2001). Exchanges of water between the main basins of the sound and the numerous fjords and bays within the sound occur in a variety of ways. As a general rule, the surface flow in a fjord is initiated by freshwater runoff at the fjord’s head. The fresh surface layer spreads along the fjord’s axis increasing its volume (by many times) and salinity through entrainment, induced by tidal currents and winds, of saltier subsurface waters. To compensate for the volume of deeper waters entrained into the outflow, subsurface waters flow inward toward the head of the fjord. This two-layer circulation, while present on average, is weak and can be easily masked in short-term observations by tidal and/or wind-driven flows. Although this generic circulation pattern holds for most fjords, there are substantial variations among fjords, owing to differences in geometry and bathymetry, winds, tides, and the salinity difference between the head and the mouth of the fjord. For example, in some seasons, the water at the mouth of the fjord might be fresher than that at the head (due to seasonal changes in river runoff into the fjord). In these situations, a reverse fjord circulation pattern can develop (Klinck et al., 1981). The renewal of deep fjord waters depends on fjord bathymetry and, in many cases, the depth of the fjord’s sill. Sills, which are usually located near the fjord’s mouth, restrict communication between the deep waters of the fjord’s inner basin and offshore waters. Deep-water renewal depends critically on sill geometry and ambient conditions. For some fjords, deep-water renewal may be infrequent and episodic, whereas for other fjords, complete or partial renewal occurs periodically either seasonally or, more frequently, through the interaction of stratified tidal flows with the sill. In such cases, tidal suction at the sill could draw deep offshore waters over the sill crest to resupply the inner basin of the fjord (Thompson and Golding, 1981; Thomson and Wolanksi, 1984). Strong tidal current– sill interactions can lead to a number of complex hydraulic effects that result in strong vertical mixing and exchange (Farmer and Smith, 1979; Freeland and Farmer, 1980).

Outer shelf and slope domain The third domain, waters over the shelf break and continental slope, includes the edge of the Alaska Current and the Alaskan Stream (west of about 150°W). These currents

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are the northern portions of the North Pacific’s subarctic gyre (discussed in the following text) connecting the Gulf of Alaska shelf and the Pacific Ocean. The Alaska Current is about 250 km wide, and has relatively weak flows (5–15 cm s−1), with correspondingly small horizontal and vertical velocity shears. In contrast, the Alaskan Stream is a narrower, 75 km wide, much swifter (50–100 cm s−1), and with highvelocity shears in the upper 500 m (Reed and Schumacher, 1987). West of 150°, the Alaskan Stream flows along the continental slope south of Alaska Peninsula and Aleutian Islands and gradually weakens west of 180°W (Thomson, 1972). The transformation of the Alaska Current into the Alaskan Stream includes changes in the velocity and density gradients along and across the shelfbreak. These gradients, in conjunction with varying bottom topography, affect the exchange between the shelf and slope (Gawarkiewicz, 1991). Hence, the transition from the Alaska Current into the Alaskan Stream implies that shelfbreak exchange mechanisms are not uniform around the Gulf of Alaska shelf. The Alaskan Stream has a mean annual transport of 20 to 30 × 106 m3 s−1 (Reed and Schumacher, 1987; Musgrave et al., 1992; Ohtani et al., 1997; Reed and Stabeno, 1999), with the broad range reflecting the uncertainty of the relatively few direct measurements. Although seasonal transport variations are apparently small (Tabata, 1991), Thomson et al. (1990) find that the Alaska Current tends to be more concentrated along the continental margin in winter than in summer. Year-to-year transport variation can be as much as 30% (Royer, 1981b) of the mean. Surface salinities vary by only about 0.5 throughout the year, whereas the magnitude of the annual seasurface temperature cycle is comparable to that of the shelf (i.e., ~10°C). Nevertheless, except for the shallow summer thermocline, horizontal and vertical density gradients are controlled by the salinity distribution. Maximal stratification occurs between 100 and 300 m depth and is associated with the permanent halocline of the Gulf of Alaska. Halocline salinities range from 33–33.8, and temperatures range from 4–6°C (Tully and Barber, 1960; Dodimead et al., 1963; Reid, 1965; Favorite et al., 1976; Reed, 1984; Musgrave et al., 1992). The upper halocline, with salinities of 33–33.3 and temperatures of about 5°C, contributes to the deep waters that flood the shelf bottom each summer (Fig. 2.12). Although flow in the Alaskan Stream appears relatively steady (Royer; 1981b; Reed and Schumacher, 1987), large (150-km diameter), clockwise eddies, moving from the interior basin onto the slope and shelfbreak, can occasionally alter the circulation here (Musgrave et al., 1992; Crawford et al., 2000; Okkonen et al., 2003). These changes include flow reversals, deflection of the shelfbreak front, and vertical displacement of the pycnocline (Musgrave et al., 1992; Okkonen et al., 2003) during eddy passage. Upwelling occurs at the leading and trailing edges of the eddy, while the pycnocline is deflected downward as the eddy center passes. The eddies are long lived (2–3 years) and support a clockwise flow with speeds between 20–50 cm s−1 (Tabata, 1982; Musgrave et al., 1992; Okkonen, 1992; Crawford et al., 2000;


Okkonen et al., 2003). They form in winter in the eastern Gulf when wind stress is strong along the eastern boundary (Willmott and Mysak, 1980; Melsom et al., 1999; Meyers and Basu, 1999) and then propagate westward at about 2–10 cm s−1. Most of the eddies remain over the deep basin and far from the continental slope (Crawford et al., 2000); however, at least one usually forms in January or February in the northeast Gulf of Alaska. Once formed, these eddies tend to propagate westward along the continental slope for several months until they approach Kodiak Island, whereupon they translate southwestward (Okkonen et al., 2003). Eddies that impinge on the continental slope could substantially influence the shelf circulation and exchanges between the shelf and slope of salt, nutrients, plankton, fish eggs, and larvae in much the same manner as the Gulf Stream rings along the East Coast of North America (Houghton et al., 1986; Ramp et al., 1983; Joyce, et al., 1992). Indeed, in the eastern Gulf of Alaska, Whitney and Robert (2002) find that these eddies cause a net transport of nutrients from the shelf to the basin.

2.2.3. Tides The Gulf of Alaska has a mixed-type tidal regime with a dominant semi-diurnal M2 tide, the diurnal K1 tide usually being secondary in importance. Tidal amplitudes and velocities are strongly influenced by the complex bottom bathymetry and geometry of the shelf and coast. Consequently, there are large tidal differences across the Gulf coast. For example, Anchorage has the largest tidal range in the northern Gulf, the M2 and K1 tides being about 3.6 and 0.7 m, respectively. At Kodiak and Seward, the tides are about half as large. The cross-shelf flux of tidal energy onto the northwest Gulf shelf is very large and is dissipated by correspondingly high frictional forces (Foreman et al., 2000). The tidal dissipation rate in Kennedy Entrance is about half of the total dissipation of the M2 constituent in the Gulf, and nearly onethird of the energy of the K1 tide is dissipated within Cook Inlet. The tidal energy dissipation mixes the water vertically, bringing nutrient-rich water into the euphotic zone. As the tidal wave passes over the rough sea bottom, it can also generate diurnal shelf waves and residual or steady flows that can be locally important for transporting suspended and dissolved materials. Diurnal-period shelf waves are prominent features along the British Columbian shelf (Crawford, 1984; Crawford and Thomson, 1984; Flather, 1988; Foreman and Thomson, 1997; Cummins and Oey, 2000), where they displace the pycnocline and alter mixed layer depths. Foreman et al.’s (2000) model results indicate that diurnal shelf waves occur in the northwest Gulf and especially along the shelfbreak east of Kodiak Island. Water stratification facilitates vertical redistribution of tidal energy over the shelf by generating internal waves at the tidal period. (Internal waves move along density discontinuities and exist only in a stratified water column. They may not be noticeable

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on the ocean surface, but the deeper wave can displace the pycnocline several meters.) Significant internal tides are likely generated at the shelfbreak in summer and fall when stratification is strong. The internal waves do not travel far (tens of kilometers), in contrast to the large scale (thousands of kilometers) of the generating tidal wave. Internal tide phases and amplitudes change with depth and seasonally with extent of stratification. While there have been no systematic studies of internal tides on the Gulf of Alaska shelf, Danielson (unpub. data) finds that the M2 tidal velocities in the ACC offshore of Seward in summer are approximately 20 cm s−1 at depths shallower than 30 m and about 5 cm s−1 below 100 m. In contrast, tidal velocities in winter are uniform over the water column at about 5 cm s−1. Internal tides can also displace the pycnocline with significant biological consequences, such as pumping nutrients into the euphotic zone, dispersing plankton and small fishes, and forming transitory and small-scale fronts that affect feeding behaviors (Mann and Lazier, 1996). Internal waves can also “break”, resulting in vertical mixing, a reduction in stratification, and the transport of nutrients into the euphotic zone. Regions of large bathymetric and tidal current gradients, such as found in Cook Inlet, often result in the formation of tidal fronts due to differential tidal mixing. These fronts typically parallel isobaths and delineate the boundary between waters that are well-mixed by tides from stratified regions. Tidal fronts can be very narrow (a few meters to hundreds of meters), and vary in strength and position over the fortnightly tidal cycle or seasonally in conjunction with changes in stratification. Rectified tidal currents and tidal fronts can also form around the perimeter of submarine banks. These regions are often highly productive and important feeding areas for fish. Although the classic example is Georges Bank off the New England coast (Horne et al., 1989), it seems likely that Portlock Bank northeast of Kodiak and Alsek Bank offshore of Yakutat behave similarly.

2.2.4. Gulf of Alaska Basin The circulation in the central Gulf of Alaska consists of the counterclockwise flow of the Alaska Gyre, which is part of the more extensive subarctic gyre of the North Pacific Ocean. The center of the Alaska Gyre is at about 53°N, and between 145 and 150°W. The gyre includes the Alaska Current and Stream and the eastward-flowing North Pacific Current along the southern boundary of the Gulf. Although some water from the Alaskan Stream apparently recirculates into the North Pacific Current, the strength and location of this recirculation is poorly understood and appears variable (Favorite et al., 1976; Emery et al., 1985). The gyral circulation pattern (illustrated schematically in Fig. 2.13) is, in fact, a three-dimensional circulation field that arises in response to the large-scale mean counterclockwise wind-stress distribution over the Gulf of Alaska. The divergence in the Ekman transports resulting from the


Figure 2.13: The major features of the Gulf of Alaska Gyre, including upwelling at mid-gyre, Ekman flow away from the center, and the currents at the shelf edge.

wind-stress distribution leads to upwelling at the center of the gyre and downwelling along the gyre boundaries. The long-term average upwelling rate is 10–30 m year−1 in the gyre’s center (Xie and Hsieh, 1995) and results in the permanent halocline being shallower here than around the edges of the gyre. It also provides a mechanism by which nutrient-rich deep waters are brought into the euphotic zone. Mean current speeds in the upper 150 m of the gyre (and far from the continental slope) are 2 to 10 cm s−1, although the variability is large (Thomson et al., 1990). Following Tully and Barber (1960) and Dodimead et al. (1963), the vertical temperature and salinity structure of the Alaska Gyre consists of: (1) a seasonally varying upper layer extending from the surface to about 100 m depth, (2) the permanent halocline between about 100 and 250 m, over which salinity increases from 33 to 33.8 and temperatures decrease from 6 to 4°C, and (3) a deep layer, extending from the bottom of the halocline to about 1000 m depth, over which salinity increases more slowly to about 34.4 and temperatures decrease from 4 to 3°C. Below 1000 m depth, salinity increases even more slowly to its maximum value of about 34.7 at the bottom. Seasonal changes in wind mixing and heat exchange with the atmosphere alter the properties of the upper layer. From October through March, cooling and strong wind-mixing establish a deep mixed layer with uniform temperature and salinity that extends to the top of the halocline. Winter mixed layer salinities range from 32.5 to 32.8, and temperatures range from 3.5 to 6°C, with the colder and fresher values

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being in the northern Gulf. As wind speeds decrease and solar heating increases in spring, the upper layer freshens slightly and warms, ultimately forming a weak halocline and a strong thermocline at about 25 m. This seasonal pycnocline erodes, and upper layer properties revert to winter conditions as fall cooling and wind mixing ensues. The halocline is a permanent feature of the subarctic North Pacific Ocean and represents the deepest limit over which winter mixing penetrates. The halocline is a consequence of both the large-scale circulation and mixing process occurring over the North Pacific and the excess of precipitation and runoff over evaporation over the high-latitude North Pacific Ocean (Reid, 1965; Warren, 1983; Van Scoy et al., 1991). Although the halocline impedes vertical exchange between the upper layer and the nutrient-rich deep water below, it is relatively shallow compared to the main thermocline of the subtropical North Pacific. Consequently, sufficient exchange, driven by storm winds, occurs even in summer to maintain relatively high nutrient concentrations in the euphotic zone. The deep waters of the Gulf of Alaska also contain North Pacific intermediate water (formed in the northwestern Pacific Ocean) and, at greater depths, include contributions from the North Atlantic. Mean flows in the deep interior are feeble (1 cm s−1), with the flow dynamics governed by both the climatological wind stress distribution (Koblinsky et al., 1989) and the global-scale thermohaline circulation (Warren and Owens, 1985), subject to modifications by bottom topography. The global thermohaline circulation carries nutrient-rich waters into the North Pacific and forces a weak and deep upwelling throughout the region (Stommel and Arons, 1960a, 1960b; Reid, 1981; Broecker, 1991).

2.2.5. North Pacific Ocean The climate and climate variability at interannual and longer time scales of the Gulf of Alaska is linked to the North Pacific atmosphere–ocean system. Although this variability is addressed in Chapter 4, this section briefly outlines the main features of the North Pacific relevant to that subsequent discussion. The global atmosphere – ocean circulation system arises in response to the imbalance in radiative heating and cooling between the equator and the poles and the asymmetrical distribution of land and water over the globe. On annual average, this disparity is erased by heat transports from low to high latitudes provided by quasiorganized atmospheric and oceanic circulation systems. The term “quasi-organized” implies that these systems have both an identifiable mean structure (that includes seasonal variations) and some degree of randomness or unpredictability, which is an integral part of both fluids. The atmosphere and ocean are dynamically connected to one another through exchanges of heat, mass, and momentum. Hence, both fluids adjust to these exchanges, and perturbations in one part of the system can (and often do) propagate to another part. For example, regional sea surface temperatures and


temperature gradients are controlled by the amount of heat in the upper ocean and the transports and trajectories of the large-scale currents. Changes in upper-ocean heat occur by heating and cooling processes, primarily by exchanges with the atmosphere, but also through mixing and advection in the ocean, along the current’s path. The volume transport and pathway of currents are mainly determined by the large-scale wind stress pattern. On the other hand, changes in ocean temperatures and patterns of air–sea heat exchange ultimately affect atmospheric pressure patterns and the wind systems over the ocean. The spatial and temporal scales over which these adjustments occur depend on the scales, magnitude, and source of the perturbation, however. In some cases, these adjustments are immediate and obvious, while in other cases, the response might take years to decades and evolve along complex pathways. The principal atmospheric features of the North Pacific Ocean are associated with the sequence of low- and high-pressure systems distributed between the equator and the pole. These pressure cells are statistical composites based on temporal and spatial averages of the many individual pressure systems that build, translate, and decay over the North Pacific. High pressure prevails over the subtropical northeast Pacific (15–35°N; the East Pacific High) and polar regions (north of 65°N), while low pressure occurs along equatorial latitudes and between 40 and 60°N (the latter being the Aleutian Low, discussed earlier). The winds associated with these pressure systems are primarily zonal (east–west) and form alternating bands of surface easterlies (westward winds) over the equator (the Southeast Trade Winds), the subtropics (the Northeast Trade winds), and poles. Westerlies (eastward winds) prevail over the mid-latitudes with the jet stream centered at about 40°N and sandwiched between the East Pacific High and the Aleutian Low. The meridional (north–south) distribution of zonal wind stress is primarily responsible for the current structure in the upper 1000 m of the ocean. In the North Pacific, the wind stress distribution divides the basin into two massive counter-rotating gyres, with each gyre containing a distinct assemblage of currents (Fig. 2.14). The clockwise subtropical gyre encompasses the width of the Pacific Ocean and extends from 10–40°N. Its southern limb consists of the North Equatorial Current, which flows westward between 10 and 20°N across the breadth of Pacific. A portion of this current turns northward along the Asian continent, where it then narrows and intensifies to form the Kuroshio Current. The Kuroshio transports approximately 25 × 106 m3 s−1 of water from the tropics to mid-latitudes and is the major oceanic pathway by which warm, salty, tropical waters are transported northward in the North Pacific Ocean. The transport increases as the Kuroshio flows northward and, at about 40°N, it turns offshore from the Asian coast to form the Kuroshio–Oyashio Extension. Here, the transport increases to about 80 × 106 m3 s−1 from the addition of water flowing southward in the Oyashio Current. This mixture continues eastward between 40 and 50°N, and east of 170°E is called the North Pacific Current. West of North America, the North Pacific Current bifurcates, with part flowing northward into the Gulf of

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Figure 2.14: The major surface circulation features of the North Pacific Ocean.

Alaska and the remainder turning southward in the California Current along the west coast of North America. The bifurcation latitude varies from about 45°N in winter to about 50°N in summer. Eventually, the California Current turns eastward to rejoin the North Equatorial Current, thereby completing the subtropical gyre. The equatorial current system consists of the westward flowing South Equatorial Current, which straddles the equator and the North Equatorial Countercurrent, which flows eastward between the North and South Equatorial currents. On annual average, the ocean gains heat south of about 30°N and loses heat to the atmosphere north of this latitude so, in aggregate, these currents affect a net northward transport of heat from the tropics. The oceanic sources of the atmospheric moisture that eventually enters the Gulf of Alaska vary seasonally (Emile-Geay et al., 2002). From spring through fall, most of this moisture is absorbed from the western tropical Pacific and the Indonesian seas and thence transported northward by the low-level winds of the Asian Monsoon. In winter, cold, dry winds spilling off the Asian mainland suck moisture into the atmosphere along the path of the Kuroshio and the Kuroshio–Oyashio Extension. Much of this latent heat exchange is controlled by the properties of the winter air masses flowing across Asia, suggesting that the strength and position of polar high-pressure systems possibly influence this process. The Subarctic Gyre consists of two small, counterclockwise-flowing gyres; the Alaskan Gyre (discussed earlier) and the Western Subarctic Gyre, which includes the Bering Sea basin. The gyres are linked to one another through the Alaskan Stream. Some of the Alaskan Stream enters the Bering Sea through the deeper passes of the


Aleutian Islands, then flows counterclockwise around the Bering Sea before leaving via a southward flow along the Kamchatka Peninsula. This outflow commingles with waters from the Sea of Okhotsk and the remnants of the Alaskan Stream to form the southward-flowing Oyashio Current. The Oyashio closes the Western Subarctic Gyre, for it feeds the Kuroshio–Oyashio Extension and eventually rejoins the North Pacific Current. The junction between the cool, fresh subarctic waters and the warm, salty subtropical waters that parallels the path of the Kuroshio–Oyashio Extension and the North Pacific Current forms the subarctic front. The warm waters, transported northward by the Kuroshio and the sea surface temperature contrast across the subarctic front, particularly within the Kuroshio–Oyashio Extension, where the sea surface temperature gradient is greatest, play a prominent role in the formation of the low-pressure systems that eventually enter the Gulf of Alaska. Most of these storms are generated in this region throughout the year, although cyclogenesis is more frequent and vigorous in fall and winter, when cold, dry air flows off Asia and encounters warmer ocean waters. Gradients in sea surface temperature result in similar gradients in sensible and latent heat fluxes, which can enhance cyclogenesis. Storms evolve as sea level pressures fall within the heated air masses and counterclockwise winds develop about the zone of falling pressure. Once formed, these storms intensify en route to the Gulf of Alaska as they continue to extract heat and moisture from the ocean (Roden, 1970).

2.3. The Marine Production Cycle Theodore Cooney 2.3.1. Introduction Photosynthesis establishes the base of the food web in most marine communities (Fig. 2.15). In the pelagic realm, phytoplankton and some bacteria are the dominant photosynthetic producers. Given sufficient light energy and nutrients (dissolved inorganic forms of nitrogen, phosphorus, silicon, and iron), these single-celled primary producers reduce carbon in the presence of sunlight and synthesize organic matter to form living plant populations. These populations are in turn grazed by a host of small animals – the zooplankton – ranging from tiny single-celled ciliate protozoans and flagellates to larger animals such as copepods and euphausiids (krill) that in turn become food for fishes, birds, and mammals. At the same time, detritus (primarily fecal material and the remains of dead plants and animals) is oxidized by bacteria as it falls toward the seabed, liberating nitrogen, phosphorus, silicon, iron, and other substances back into the water. Cycling of inorganic nutrients and carbon to living organic matter (a synthesis process) and back to dissolved inorganic forms (a remineralization process) defines the marine production cycle at lower trophic levels.

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Figure 2.15: Organic matter cycling in the pelagic ocean. See text for description.

Unlike terrestrial ecosystems that are rich in stored carbohydrate and are relatively slow growing, the pelagic marine system is dominated by a protein ecology with little long-term energy storage but very rapid growth (Parsons, 1976). A hallmark of this system is that there is little carryover of plant matter from year to year.

2.3.2. The Annual Cycle of Production The rate at which organic matter is produced in the ocean depends on the cyclic availability of both nutrients and light (see Fig. 2.20). In mid-to-high latitudes, there is only enough daylight for significant phytoplankton growth during a part of the year – spring and summer into early fall. Within this period, plant growth can continue until nutrients are exhausted. During the dark days of winter, production falls to very low levels. At the same time, cooling of the surface waters and increasing winter storm activity deepens the wind-mixed layer, causing dissolved inorganic nutrients from deep sources to become entrained and brought to the surface. In this


BOX 2.1: PLANKTON by Theodore Cooney By definition, plankton is the community of tiny marine plants and animals whose distribution is influenced primarily by ocean currents. Phytoplankton (the photosynthetic plants and bacteria) (Fig. 2.16), zooplankton (Figs. 2.17 and 2.18) (the small drifting animals), and meroplankton (eggs and larval stages of larger organisms) in this community are generally most abundant near the surface of the ocean – the upper 200 m. Although their distributions are at the mercy of the currents, some zooplankters undergo extensive vertical migrations (up to several hundred meters) in response to seasonal cycles in their food abundance and/or reproductive strategies. Also, many zooplankters migrate daily – swimming up into the surface waters at night, where they feed or mate, and returning to depth during the day to digest their food and escape predators. The majority of marine plants are unicellular and can double their populations approximately every day under ideal conditions. In the seasonal ocean, the living mass of plankton varies over an order of magnitude or more each year, usually reaching a peak in the summer before falling to low levels in mid-to-late winter. The dramatic increase in biomass that begins early each spring is called the plankton bloom. Marine scientists use fine-mesh nets and water bottles to sample plankton (Fig. 2.19).

Figure 2.16: Phytoplankton: diatoms (photograph courtesy of NOAA).

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Figure 2.17: The large pelagic copepod Neocalanus cristatus (photograph courtesy of R. Hopcroft, University of Alaska).

Figure 2.18: The pteropod Limacina sp., a pelagic mollusk (photograph courtesy of R. Hopcroft, University of Alaska).


Figure 2.19: Sampling with a bongo net (photograph courtesy of EVOS Trustee Council).

way, upper-layer nutrient concentrations are replenished each year prior to the next spring bloom (see Section 2.2). Sverdrup (1953) developed an elegant theory that explains the timing of onset of the plankton bloom in the spring. In the presence of sufficient nutrients, phytoplankton growth rates are primarily determined by light intensity. Since light is attenuated logarithmically with depth, a sinking cell will quickly reach a depth where available light permits just enough photosynthesis over 24 h to balance respiration – the compensation light intensity. This depth generally defines the seasonally changing base of the photic zone. Cells sinking below this depth receive less than the compensation light intensity and cease to grow. The critical depth is the depth above which the overall depth-integrated photosynthesis just equals the depth-integrated respiration over a 24-h period for a population of plant cells distributed uniformly by wind mixing. From this depth to the surface, the mixed-layer phytoplankton population experiences an average 24-h light intensity that just balances its respiration. When the mixed layer extends below the critical depth, cells in the population receive insufficient light to balance their respiration, and production ceases. Prediction of the bloom timing follows from these limiting conditions. In the winter, low surface light and deep vertical mixing establishes a mixed layer that is much deeper than the critical depth (Fig. 2.20). As a result, the rate of primary

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Figure 2.20: The annual cycle of primary productivity including water, light, wind, and nutrient conditions that initiate the spring bloom, support periodic spring and summer production, and lead to the end of the production, season in the coastal Gulf of Alaska.

productivity – photosynthesis – is extremely low. As days lengthen in the spring, increasing sunlight extends the critical depth deeper into the water. At the same time, decreasing winds and fresher, warmer surface water cause the mixed layer to contract towards the surface. When the mixed layer becomes shallower than the critical depth, all plant cells in the mixed layer receive sufficient light energy to exceed their respiratory demands. The result is an immediate expansion of phytoplankton biomass, now growing rapidly on high concentrations of dissolved nutrients – the spring bloom. Because the timing of this explosive growth depends on weather-forced vertical mixing interacting with seasonally available light, the conditions for this event rarely occur on the same date each year, and the start of the bloom typically varies by a week or more between years. The main variables that appear to be responsible for the variation in timing are the degree of vertical stability determined mainly by winds and freshwater input, the amount of cloud cover, and the angle of incident radiation.


Once underway, the annual spring burst of primary production continues as long as adequate nutrients are available. This state is eventually terminated by increasing stratification as a warmer, less saline surface layer develops over a body of deeper, cooler, more dense water, dramatically reducing vertical exchange and nutrient renewal to the upper layers. Some production continues after this time, due to (1) episodic wind-mixing events that reinject nutrients into the surface layer and boost growth for short periods (see wind events in Fig. 2.20), (2) bacterial oxidation of detritus which recycles some inorganic nitrogen, mostly in the form of ammonia, and (3) nitrogen fixation by the cyanobacteria. Overall, primary production declines to low levels following stratification, and the resulting seasonal hiatus in plant growth is termed the nutrient-limited portion of the production cycle. In some locations where nutrients are continuously available in the surface waters (for example, through persistent upwelling from deep sources as occurs in the entrance to Cook Inlet, or in the open ocean), phytoplankton production can continue in proportion to light levels throughout the summer. In summary, over the course of the year, the general seasonal pattern in photosynthesis includes a period of severe light limitation during the winter, short episodes (a few weeks) of dramatic growth in the early spring (and possibly the fall), and nutrient limitation during the late spring, summer, and early fall. The examples of these patterns from the Gulf of Alaska and their regional exceptions will be discussed in the following text. The explosive spring bloom in plant plankton is followed quickly by a similar expansion in numbers of the various zooplankers that graze on it (Fig. 2.21). In this way their seasonal populations parallel that of the phytoplankton, but with varying time lags, depending on the species. For example, the small protozoans can keep pace reproductively with the plants, so their populations fluctuate in close synchrony. In contrast, the larger zooplankters take longer to reproduce – weeks or months – and most must first feed before they can produce offspring. The result is a time lag between the peak of the phytoplankton production and seasonal peaks in the large zooplankton community, which generally come later in the growing season. As the biomass of plant and animal plankton expands, the dominant species follow a predictable progression. At the beginning of the season, diatoms dominate in the cold, high-nutrient waters. These plants, many of them producing large colonies, have an external covering of silica and are often quite large – some cells greater than 50 microns in size. When nutrient concentrations decline in the late spring and early summer, the phytoplankton community shifts to smaller species with increased surface-area-to-volume ratios. Because nutrients are absorbed through the cell walls, increasing surface area relative to cell volume maximizes nutrient uptake for the plants. The result is that nutrient-poor summer waters typically host very small phytoplankton – some tiny diatoms (5 microns and smaller), and a larger proportion

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Figure 2.21: The progression of various plankton populations, their trophic linkages, and the availability of nutrients during the spring and summer production season.

of microflagellates and medium-sized dinoflagellates. In addition to being small, the flagellated forms swim and can thus exploit more optimum levels of light and nutrients. The seasonal progression in size of the primary producers is mirrored by similar shifts in the zooplankton. The early-spring zooplankton community is dominated by large calanoid copepods and euphausiids – primarily filter feeders – which target the diatom populations. Some of these consumers (and sometimes their progeny) overwinter in the deep water and migrate back to the upper layers to feed in the spring. When the phytoplankton shifts toward smaller forms in response to nutrient exhaustion, the larger copepods and euphausiids are replaced, in part, by species that are more capable of feeding on particles 5 microns in size and smaller. During this time, the Pteropoda (mucus-feeding pelagic molluscs) and the Larvacea (very fine, filter-feeding pelagic tunicates) expand their populations. Also, during the summer and fall, carnivorous jelly plankton (Ctenophora and Cnidaria) become well established as the community shifts away from domination by herbivores toward a greater diversity in feeding types – more omnivores and carnivores.


The fundamental processes described in the preceding text form the basis of the marine production cycles in the northern Gulf of Alaska. However, there are some important exceptions and regional differences. In the following sections, we will compare plankton production cycles in three distinct areas: (1) protected inner waters, (2) open ocean, and (3) shelf and coastal waters. Figure 2.22 provides a comparison of important features of the production system in these domains.

Protected Inner Waters The production cycle of protected inner waters (sounds, fjords, and inlets) follows the classical picture described earlier. There is a well-developed spring burst of plant production and biomass accumulation (see Fig. 2.23) that, at first, features populations of large diatoms and then gradually gives way to smaller photosynthetic forms as the water column stratifies and nutrient supplies dwindle in the photic zone. This general pattern can be illustrated by measured changes in water fluorescence (a proxy measurement for chlorophyll a, the photosynthetic pigment for most algae) for Prince William Sound (Eslinger et al., 2001). Continuous measurements at a depth of 10 m over the year captured the diatom bloom that typically peaks in mid-April, followed by a rapid decline to summertime lows beginning in May and June. A second fall “event” is documented for October and November (Fig. 2.20). Similar patterns in

Figure 2.22: The contrast between inner- and outer-shelf plankton community structure, production, and how inshore–offshore differences in trophic efficiencies arise.

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Figure 2.23: A false-color satellite image of the northern Gulf of Alaska taken on April 5, 2002, showing the high concentrations of chlorophyll a in the surface waters of Prince William Sound and some other isolated inshore and offshore locations. Red indicates the highest concentrations of chlorophyll a, and blue and green colors indicate the lowest (image courtesy of D. Musgrave and R. Potter, University of Alaska; data from NASA).

plant pigments have been reported for at least the late winter through summer periods in coastal Alaska (Goering et al., 1973; Iverson et al., 1974; Burrell, 1986; Larrance et al., 1977). Annual primary productivity estimates for protected waters range from 300 g C m−2 for lower Cook Inlet to 185 g C m−2 in northwestern Prince William Sound and 145 g C m−2 in Boca de Quadra in southeastern Alaska. These values are similar to those of protected marine waters in Norway, Sweden, Greenland, and Canada (Cooney and Coyle, 1988). Studies of nutrient availability in the upper layers of protected coastal waters document nitrogen enrichment during the light-limited period of the production cycle in winter. Nitrate and silicate concentrations decline rapidly after the initiation of


the spring boom (Goering et al., 1973; Heggie et al., 1977; Burrell, 1984; Ward, 1997), low levels of primary productivity being sustained through the summer by nutrient recycling in the photic zone (ammonia production), by nitrogen fixation, and by localized upwelling and vertical turbulence (Sambrotto and Lorenzen, 1987). During the period of relaxed coastal downwelling in each summer (more so in some years than in others), the intrusion of subsurface slope waters onto the shelf provides a mechanism to enrich the deeper coastal regions with plant nutrients. These nutrients find their way into the productive surface layers through deep convective and wind mixing in the following winter and early spring, and also by other processes such as tidal pumping. The phytoplankton of sounds and fjords is dominated by large diatoms during the spring bloom (Ward, 1997; Horner et al., 1973), while the summer plant assemblages are primarily smaller forms. Chaetoceros, Thalassiosira, and Skeletonema are among the most common large diatom genera, although their dominance ranking typically changes from year to year. The zooplankton in protected inner waters is a combination of oceanic and neritic forms. In the early spring, zooplankton community biomass in Prince William Sound is dominated by the large calanoid copepods Neocalanus spp., Calanus marshallae, and Metridia okhotensis, with barnacle nauplii and the early copepodites (C1 and C2) of Neocalanus being important contributors as well (Cooney et al., 2001). By early summer, the community biomass switches to dominance by the small calanoid, Pseudocalanus spp., the pterpod, Limacina helicina, and the larvacean, Oikopleura sp. Later in the summer, in the freshened and warm upper layers, Pseudocalanus, ctenophores, and arrow worms, Sagitta elegans, are dominant in most samples. Wetweight biomass in the upper 50 m averages about 10 mg m−3 in the winter and early spring but peaks near 600 mg m−3 in midsummer – late June and July. Copepods dominate the numbers and biomass of most net-caught samples during all seasons.

The Open Ocean The open ocean ecosystem differs remarkably from the protected inner waters. First, there is little or no observable spring bloom. Phytoplankton numbers generally remain low throughout the year; the only hint of a plant bloom may occur occasionally in the fall, not the spring; and the seasonal accumulation of plankton occurs as zooplankton, not phytoplankton. Second, inorganic nitrate, a principal limiting nutrient in shelf and inner waters, is only rarely exhausted in the surface waters of the open ocean. This apparent puzzle, i.e., abundant nitrate yet no phytoplankton bloom, has prompted a number of ideas about how the open ocean functions. Much of the story has been put together over the last 25 years for the North Pacific Ocean (Miller, 1993). Trophic structure and nutrient dynamics in the open ocean of the Gulf of Alaska are very similar to oceanic ecosystems elsewhere in the world that are termed

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high-nutrient, low-chlorophyll (HNLC) areas. Compared to the classical picture observed in protected inner waters, there are four primary differences in these areas: (1) the main plants at the base of the food web are very small diatoms and flagellates (2–10 microns); (2) the primary grazers are small protozoans; (3) the dominant source of nitrogen for plant growth appears to be ammonia, not nitrate; and (4) in all HNLC areas investigated, plant growth has been shown to be iron limited. The lack of an observable spring phytoplankton bloom is probably due to the presence of very efficient grazers that eat the phytoplankton as quickly as the latter can grow and divide, even during the optimal conditions in the spring. So, the rapid spring growth of microflagellates responding to increasing day length is kept pace with by grazing ciliates and other protozoans. Because deep vertical winter mixing in the oceanic subarctic Pacific is restricted by a permanent halocline (steep gradient in salinity) at 100 m, plants and micrograzers are rarely diluted to levels that would cause a trophic uncoupling of the plants and tiny herbivores. Since the protozoans are capable of equaling the reproductive rates of the plants, there is rarely, if ever, a time during the year when the primary producers are free from grazing losses and capable of expanding their populations. The large calanoid copepods are important grazers in the open-ocean plankton communities, as in inner waters, but occupy a different trophic position. These copepods exploit the epipelagic zone (upper 200 m) for feeding and growth, but use the deeper mesopelagic zone (500–1000 m) for diapause and reproduction. In contrast to inshore waters, where they can flourish feeding on large phytoplankton, the large copepods cannot get enough energy from eating small primary producers (Dagg, 1993). Instead, they graze on the protozoans, which are the primary grazers. The year-round presence of nitrate in the photic zone of the open ocean may be partially explained by differences in nitrogen cycling. On the continental shelf, nitrate is the dominant form of nitrogen utilized by plants, but ammonia is much more important in the open ocean (Wheeler and Kokkinakis, 1990). Ammonia is the first oxidative product of the remineralization process, and because it is relatively high in energy, it is “preferred” by phytoplankton over nitrite and nitrate. Ammonia is produced by bacterial oxidation of organic matter, and also excreted by zooplankton (and other consumers). The highly developed protozoan community of the open ocean (along with other consumers) probably supplies large quantities of ammonia to the surface waters. As long as ammonia is present in the photic zone, nitrate uptake is negligible. Because nitrate, which ultimately comes from deep wintertime mixing and weak upwelling in the central gyre is not appreciably utilized, the remainder is presumably available for transport onto the shelf for utilization there by the diatomdominated bloom. Another potentially contributing factor to the nature of the offshore pelagic community is iron limitation. Iron is probably supplied to the surface waters of openocean regions around the world by dust blown in from surrounding land, sometimes


over great distances. This essential plant micronutrient greatly affects oceanic primary productivity in the subarctic Pacific (Martin and Fitzwater, 1988). There are several open-ocean regions that exhibit high nutrients and low chlorophyll (HNLC) conditions besides the subarctic Pacific, and all have demonstrated an increase in primary productivity when enriched with iron. Because of these findings, some have suggested that the subarctic Pacific might owe its characteristic small-sized pelagic flora and seasonally residual nitrate to oceanic iron deficiency (Miller, 1993). Thus, while some questions remain, the most recent studies suggest that, in the open ocean, the sustained nitrate levels and the dominance of small-sized plant cells is most likely due to a combination of iron limitation and ammonia production (Strom et al., 2000). In summary, the seasonal picture that emerges from plankton studies in the open ocean of the subarctic Pacific begins with light limitation in the winter and early spring. Winter mixing and central gyre upwelling enrich the surface layers with dissolved inorganic nutrients, while photosynthesis is at a minimum. With the advent of weak stratification in the spring (associated with a warming water column), photosynthesis begins in the plant communities. However, because the micrograzers can apparently crop the phytoplankton very efficiently, little or no accumulation of plant biomass occurs, and it is the grazing community that increases in stock size. This increase is particularly evident for the large copepods (easily sampled with nets) – Neocalanus spp. and other large forms. As the season progresses, inorganic nutrients are drawn down in the photic zone, but the rapid recycling of ammonia production by bacteria and other consumers, along with what appears to be a dissolved iron deficiency, seemingly prevents the nitrate levels from dropping to zero. Annual primary productivity for the oceanic subarctic Pacific is thought to exceed 100 g C m−2 (Welschmeyer et al., 1993). With the renewal of fall and winter cooling, wind-mixing, and decreasing light levels, the production cycle again enters its light-limited winter phase. Late fall, winter, and early spring domination of the northeastern Gulf of Alaska by the Aleutian low-pressure system produces a shoreward Ekman flow and coastal convergence for more than half the year. Although the shoreward transport is relatively weak (approximately 5 km day−1) and pulsed by storms, it is apparently sufficient, when operating over several months, to introduce sizable quantities of biomass of oceanic origin to shelf and coastal ecosystems (Cooney, 1988; Kline, 1999). How much of this material is imported is unknown, although some believe that it significantly enriches coastal and shelf food webs in some years (Kline, 1997).

Open Coastal and Shelf Waters Production in the open coastal and shelf waters is not as well characterized as the protected inner waters and the open ocean. There are only a few studies of rates of primary production, but this environment is now the focus of studies that should become available soon. Sambrotto and Lorenzen (1987) provide a thorough historical

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review of phytoplankton and seasonal primary production for the Gulf of Alaska, including open coastal and shelf waters. The largely speculative estimates of KoblentzMishke et al. (1970) placed the production for shelf areas at 100–150 g C m−2 year−1. However, studies by Larrance (1971) of Adak Bay and shelf waters, by Larrance and Chester (1979) of outer Cook Inlet, and by Larrance et al. (1977) of the Kenai shelf provided annual estimates of primary production that ranged much higher; 300–330 g C m−2. In the absence of developed literature, it seems reasonable to suggest that the waters of the open coast and shelf of the northern Gulf of Alaska contain plant and animal plankton that is a mixture of organisms advected shoreward from the deep ocean to the coastal convergence, and those carried seaward with the seasonal offshore spread of freshened waters in the summer and early fall. Some work carried out in this region under Outer Continental Shelf Environmental Program sponsorship in the 1970s demonstrated the seasonal intrusion of oceanic zooplankton, presumably carried shoreward under active downwelling in the fall, winter, and spring months (Cooney, 1988). In addition, the landward margin of the shelf hosts the Alaska Coastal Current, which serves to distribute and mix plankton communities over 2000 km of shoreline, from northern British Columbia to the tip of the Alaska Peninsula. New studies of the shelf in the northern Gulf of Alaska sponsored by the North Pacific GLOBEC program focus on the spatially distributed attributes of the annual production cycle – nutrient sources and sinks, the role of microplankton, organic matter transfer processes, and the distribution, timing, and composition of plankton blooms in relation to factors influencing the survival of juvenile salmon entering this region from coastal nurseries. These studies are beginning to describe a much more complex shelf system than previously reported (Weingartner et al., 2002). Preliminary studies of the shelf break blooms seem to indicate that the dominant phytoplankton are small diatoms and that protozoans are important grazers (GLOBEC, unpublished). As more of these studies are reported, the structure and dynamics of the shelf ecosystem will emerge.

2.4. The Transfer of Matter and Energy Through the Food Web Theodore Cooney The marine production cycle in the northern Gulf of Alaska is strongly seasonal at the base of the food web, with periods of low productivity in the winter, an explosive spring bloom, but followed by often discontinuous growth in the remainder of the spring, summer, and fall months. The interplay between available light and inorganic plant nutrients – modified as we have seen by ocean physics – is primarily responsible


for the seasonal variability. The larger animals supported by seasonally shifting production at lower trophic levels must cope with a forage environment that cycles (only somewhat predictably) between feast and famine each year. The burst of production in the spring spreads quickly through the food web. For example, yellowfin sole (Pleuronectes asper), walleye pollock (Theragra chalcogramma), and Pacific cod (Gadus macrocephalus) replenish energy reserves (lost over the winter and from spawning) each spring in a short period of intensive feeding (a few weeks) (Paul and Smith, 1993; Paul 1997; Smith et al., 1988, 1990) (Fig. 2.24). Similarly, adult walleye pollock and Pacific herring feed on dense layers of near-surface macrozooplankters (primarily, large calanoid copepods, krill, and pteropods) to rebuild their post-spawning energy reserves during late April, May, and early June (Willette et al., 2001). The flow of organic matter in some parts of marine food webs is an intermittent process tied to the interlocking life histories, behaviors, and trophic requirements of predators and prey and modified by the seasonal dynamics of the production cycle. These habitat dependencies tie each species to the biological and physical vagaries of the marine environment that support them over time. As we will see, the food supply for fishes, birds, and mammals is often scattered and ephemeral, with most energy coming from a few key species.

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Figure 2.24: Seasonal changes in whole-body energy content of yellow sole Pleuronectes asper in the Gulf of Alaska (after Paul, 1997).

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2.4.1. Food-Web Structure – The Principal Forage Stocks Pelagic and benthic food webs in the Gulf of Alaska are a complex array of hundreds of species linked in a poorly understood trophic process that moves matter and energy from primary producers to all consumer levels. Approximately 300 species of fishes, 147 species of birds, and 26 species of marine mammals feed in the Gulf of Alaska for at least some portion of each year (OCSEAP, 1986; DeGange and Sanger, 1986; Calkins, 1986). While the forage utilized by apex consumers is diverse overall, relatively small collections of invertebrates and small schooling fishes are believed to provide the bulk of the food (Springer and Speckman, 1997). These trophically important or key forage stocks include a few macrozooplankters (two genera of large calanoid copepods, a pteropod, and euphausiids), juvenile herring, salmon and pollock, capelin, sand lance, eulachon, and two or three mesopelagic fishes – fewer than 20 in all (Fig. 2.25). Most of these forage species exhibit schooling, layering, and/or swarming behaviors during some or all of their life histories, providing dense feeding opportunities for consumers who can exploit these patches. In this manner,

BOX 2.2: FOOD CHAINS AND WEBS by Theodore Cooney “Food chains” and “food webs” are terms used to convey word pictures describing some aspects of consumptive processes. Energy fixed by photosynthetic plants and bacteria is transferred to consumers through “trophic” or feeding relationships. Descriptions of forage dependencies, for example, by seabirds, provides a way to conceptually understand how some parts of a consumptive system operate. Early studies of energy flow through aquatic systems (marine and freshwater) made some simplifying assumptions about the structure of communities, electing to assign “trophic levels” to various consumers. This structure began with photosynthetic “primary producers” at level 1, herbivores (primarily zooplankton) at level 2, and first and secondary consumers at successive higher levels. As the sophistication of trophic studies grew, it became apparent that the consumptive process was quite complicated and actually more web-like, with some consumers, such as fishes, occupying several different “trophic levels” during different times in their life histories. Recent advances in numerical modeling studies are using this information to begin constructing parts of the trophic process for predictive and experimental purposes.


Figure 2.25: Waist species that funnel much of the matter and energy into the predatory fish, seabirds, and mammals. Upper row, left to right: calanoid copepods, euphausid, copepod, and pteropod; middle left: sand lance; lower left: capelin; and lower right: juvenile pollock [photographs courtesy of R. Hopcroft, University of Alaska (copepods, euphausid, and pteropod), NOAA (sand lance, capelin), and the Japan Agency for Marine-Earth Science and Technology, JAMSTEC (pollock)].

the food web in the northern Gulf of Alaska is dramatically constricted at intermediate levels for many apex consumers by the dominance of these few but locally and seasonally abundant forage stocks. An example of this “waist” can be seen in seabird diets (percentage weight in stomachs) in Prince William Sound that demonstrated (in descending order) the importance of some of these forage stocks for 14 different seabirds: (1) juvenile herring 22.4%; (2) sand lance 22.0%; (3) capelin 14%; (4) macrozooplankton 6.3%; (5) juvenile pollock 6.0%; and (6) juvenile salmon 1.0%. These six categories accounted for 72% of all foods (12 categories) by weight in the stomachs of the birds (Fig. 2.26) (Fisheries Center, 1998). Aggregations of macrozooplankton provide excellent forage. The surface-swarming and layer-forming large oceanic copepods – Neocalanus spp. and Calanus marshallae – exhibit peak biomasses in the spring and early summer (Cooney et al., 2001). This occurs just before the maturing older stages (adults and C5 copepodites) leave the upper layers to enter a diapause in the deep waters (>400 m). From mid-April through early June in protected coastal areas such as Prince William Sound, copepodite stages C4 and C5 occur in dense, near-surface layers that stretch unbroken for tens of kilometers. Abundances exceeding 500 individuals m−3 are common in

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JUVENILE SALMON JUVENILE POLLOCK MACROZOOPLANKTON

JUVENILE HERRING

CAPELIN

SAND LANCE

Figure 2.26: The composition of major forage groups in seabird diets in the Gulf of Alaska (after Fisheries Center, 1998).

these layers (Kirsch et al., 2000). (Fig. 2.27) The C5 copepodites range from 3.0 to 5.0 mm in length, a size that makes them particularly vulnerable to planktivorous fishes (pollock and herring), some seabirds, and even humpback whales. The pteropod – Limacina pacifica – because of its size (5–10 mm), abundance, and swarming behavior also provides forage opportunities for near-surface feeders in late May and June (Cooney et al., 2001). In contrast to the seasonally dominant large copepods and other macrozooplankton, euphausiids – Euphausia pacifica and Thysanoessa spp. – are present near the ocean’s surface throughout the year at night (Fig. 2.28), and in extensive layers and swarms below the surface in shelf and coastal waters at about 100 to 120 m during the day. Thysanoessa spp. spawn at the surface during the spring diatom bloom in coastal and shelf waters. Adults older than about 18 months die after spawning, and their remains are sometimes washed up on area beaches. Spawning swarms provide excellent forage for whales, seals, and some birds. It is the small schooling fishes – capelin, Pacific sand lance, eulachon, Pacific herring, juvenile salmon, juvenile pollock and cod, and a few mesopelagic fishes – that are most often singled out as the dominant forage stocks in the Gulf of Alaska for larger fishes, birds, and mammals (Springer and Speckman, 1997). Their spawning aggregations often provide forage that is extremely important to their predators. Sand lance spawn in September and October near the shore on sandy or fine-gravel beaches (Robards et al., 1999a). Capelin, herring, and eulachon are all spring spawners; the former two laying eggs in the intertidal zone, while the latter is anadromous, spawing in local rivers.


Figure 2.27: A multichannel acoustic backscatter record (200 kHz) taken on the shelf off Seward, Alaska, in May 1998 at night, showing the dense surface layers of plankton. The red indicates the greatest backscatter and probably the densest aggregations of animals and the distance from shore (top) (image courtesy of K. Coyle, University of Alaska). Two calanoid copepods that contribute to these plankton spring layers, Calanus marshallae and Neocalanus cristatus (bottom left, right, respectively) (photographs courtesy of R. Hopcroft, University of Alaska).

The spawning of Pacific herring (usually in March/April) is a particularly momentous ecological event – tens of miles of shoreline turn white with milt, and portions of the intertidal zone are covered with half a meter or more of eggs (see Fig. 4.17). Birds, fishes, marine mammals, and large invertebrates gather to feed on the spawning adults and their eggs. Bishop and Green (2001), working in Prince William Sound in 1994 found that, five species of birds consumed 31% of the eggs deposited in the region that year. During a spawning event, stocks of adult herring are also targeted by

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Figure 2.28: Three common euphausids in the gulf: Thysanoessa intermis (top), Thysanoessa longipes (middle), and Euphausia pacifica (bottom) (photographs courtesy of R. Hopcroft, University of Alaska).


humpback whales and Steller sea lions. Later in the summer, age-0 juvenile herring begin to appear in protected bays, inlets, and fjords termed nursery areas (Norcross et al., 2001). From this time on, the developing and mature herring represent one of the most significant forage bases in the coastal northern Gulf of Alaska (Brown, 2003). In Fig. 2.29, schools of herring and sand lance can be seen in nearshore areas of Prince William Sound.

Figure 2.29: Aerial photographs of nearshore schools of juvenile herring (top) and sand lance (bottom). Arrow indicates long, gray, arc-shaped school of sand lance in front of vessel (photographs courtesy of E. Brown, University of Alaska).

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Little is known of the life history and seasonal distribution of capelin, an ephemeral but very important forage fish. Fished commercially in the northwestern Atlantic Ocean, this species is not exploited in the Gulf of Alaska. Capelin is a cold-water, pelagic species, inhabiting arctic and subarctic zones in both the Atlantic and Pacific. It can form very large schools in inshore waters. During the demise of the northwest Atlantic cod stocks during the 1990s, capelin increased, zooplankton decreased, and indices of phytoplankton rose (Carscadden et al., 2001). It was hypothesized that this rearrangement was the result of the significant downturn in cod population (because of over fishing and environmental causes), releasing the capelin from their major predator. Increasing capelin stocks placed greater pressure on zooplankton, which in turn grazed less phytoplankton. These changes were described as a top-down trophic cascade triggered by the dramatic declines in cod biomass during the 1990s. Spawning aggregates or capelin appear irregularly on shelf waters, sometimes in very large numbers – schools several kilometers in size. Large numbers of predatory fish, seabirds and marine mammals, including humpback whales, feed on these large capelin schools. In Fig. 2.30, thousands of seabirds, mainly shearwaters (Puffnus spp.), and many humpback whales are shown feeding on a school of capelin near the Barren Islands. Sand lance are schooling zooplanktivores that are common in the entire edge zone of the Gulf of Alaska (Springer and Speckman, 1997). They occur in abundance

Figure 2.30: Humpback whales and seabirds feeding on a very large school of capelin near the Barren Islands, lower Cook Inlet (photograph courtesy of A. Kettle, USFWS).


in eastern Prince William Sound, where a stock estimated to be around 70 tons occurs (E. Brown, pers. comm.). Sand lance aggregate in dense shallow-water schools, easily observed from low-flying aircraft. In samples of nearshore fishes taken with beach seines, sand lance were the most abundant fish taken off sandy beaches (Robards et al., 1999b). The maximum energy content of adult sand lance is reached in the spring, which then declines through the summer and into the fall. They spawn in the fall in Kachemak Bay in lower Cook Inlet (Robards et al., 1999a) and spawn relatively few yolky eggs that develop attached to the seabed just below the lowest tides. Sand lance burrow into soft-bottom sediments at night and perhaps over longer periods during the winter. Sand lance represent a major food item for a large number of seabirds in the Gulf of Alaska (DeGange and Sanger, 1986), and their presence or absence can be crucial to the reproductive success of some seabirds. Eulachon are anadromous pelagic zooplanktivores living on the outer shelf of the Gulf of Alaska but returning each spring to spawn in freshwater (Love, 1966). As a result, these smelts are only temporary residents of the coastal waters, where they are fished heavily for subsistence purposes. When they school on the inner shelf, eulachon are consumed by most of the same predators that eat capelin and herring. At some locations, juvenile pink salmon can be a substantial forage base. In Prince William Sound, the combined release of hatchery-reared juveniles and those entering naturally from streams and small rivers comes close to 730 million fish annually. Willette et al. (2001) calculated that, during the period 1994–1996, as many as 75% of these juveniles could be consumed by predators during their first 45 days in the ocean. If the juveniles enter saltwater at 0.3 g each and grow at 4% of their body weight per day, and further, if the mortality all occurs half-way through the 45-day period (for computational ease), predators eating these 547 million juveniles will conservatively consume about 400 MT of fry before the young fishes have a chance to adapt to their new environment. Local populations of seabirds, adult herring and pollock, and several juvenile gaddid, have been documented as the major juvenile pink salmon predators (Willette et al., 2001; Scheel and Hough, 1997). Finally, it is now recognized that all forage is not the same when it comes to energy content – some of the small schooling fishes and zooplankters are much more “fatty” than others. Feeding experiments comparing the growth of marine mammals and seabirds reared on foods of different energy densities have demonstrated that growth, survival, and reproductive capacity are all enhanced when the food is rich in lipids (see Springer and Speckman, 1997, for a review). Since the dominance patterns of forage stocks vary considerably over time, some “fatty” species such as herring will support enhanced consumer growth in some years, but not in others when “less-fatty” species (such as juvenile pollock) become dominant. For sand lance, their maximum energy content occurs soon after the spring plankton bloom and coincides with the breeding seasons of many of their predators, e.g., seabirds and some marine

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mammals (Robards et al., 1999a). The effects of different energy content of prey can be partially modified by consumers who are able to switch to other forage sources when their preferred food is limiting. However, in the case of some seabirds, switching to another forage base may not be an option because of foraging range and the depth of feeding limitations (see Section 2.5.3).

2.4.2. Efficient Foraging on Patches Aggregating behavior of forage species probably makes the transfer of production through the food web relatively more efficient than if these aggregations did not occur. Supporting evidence comes from early attempts to model the marine ecosystem, including commercial stocks of pelagic and demersal fishes in the Bering Sea (Laevastu and Favorite, 1981). Modelers were surprised to discover that the levels of zooplankton required to sustain observed fish production were apparently much higher than reported from field studies. This discrepancy was eventually resolved (at least in part) when it was realized that plankton net samples integrate much of the important in situ small-scale patchiness that represents critical feeding opportunities in the ocean; the statistically averaged or operationally integrated abundances are only abstractions of the “real world,” not true characterizations of the actual feeding environment. Recent advances in acoustic and optical methods

BOX 2.3: IMPORTANCE OF TIMING by Theodore Cooney In a strongly seasonal ocean such as the northern Gulf of Alaska, conditions supporting productivity at the base of the food web are constantly changing. While the overall temporal pattern of the annual production cycle is generally similar from year to year, in reality the specifics of any 2 years are never exactly the same. Some years, the plant bloom is early, intense, and truncated, while for other years, it is late, moderate, and extended. Consumers living in this system have adapted their life cycles to address a kind of “general” seasonality in critical forage availability. This means that when conditions deviate markedly from the norm, parts of the food web become disconnected or less than optimally functional, and some consumers face failing conditions. One successful strategy adopted by some fishes and invertebrates to address this temporal and spatial “match/mismatch” problem is to extend the period during which they release eggs and larval forms, thereby broadening the temporal window to ensure that at least some progeny will find optimal survival conditions.


of census demonstrate the presence and trophic importance of a “grainy” feeding environment. In the northern Gulf of Alaska, “patch-dependent feeding” is fostered by the behaviors and life histories of the dominant forage species, especially reproductive aggregations, which, as we have seen, serve to transfer immense quantities of matter and energy through selected parts of the food web (Thomas and Thorne, 2003). Patchiness is also fostered by the physical environment where fronts and eddies serve (under some conditions) as collection sites for zooplankton and other weakly swimming pelagic forage populations. In this way, food sources are spatially distributed in relation to the physics and geology, while at the same time being temporally ordered in ways that reflect the unique characteristics of the annual production cycle at each time and place. The timing of the spring diatom bloom in Prince William Sound differs from year to year by as much as three weeks (Eslinger et al., 2001). This variability probably affects the coupling or trophic phasing between producers, first-order consumers, and higher trophic levels. However, little is known about the ecological consequences of early or late blooms. Also, there has been little attention paid to the trophic implications of renewed fall productivity in coastal and shelf waters. For some fishes that depend on late-season energy provisioning prior to winter – such as juvenile Pacific herring – the fall portion of the production cycle may be critically important to the success or demise of a year-class.

2.4.3. Food-Web Complexity and Efficiency Food chain length, or the number of steps between photosynthetic producers and the apex consumers of the production, is determined primarily by the physical size of the dominant primary producers (Ryther, 1969). The surface waters of most deep ocean environments are typically composed of small phytoplankton responding to nutrient limitations (such as iron). In these environments, more trophic exchanges are needed to generate stocks of suitably sized forage organisms than in shelf or coastal regions where larger primary producers (colonial and chain-forming diatoms) feed some forage stocks directly (see Fig. 2.31). For example, large diatoms feeding krill, which then immediately serve as food for some birds and mammals, defines a highly efficient two-step transfer mechanism. Conversely, tiny diatoms and flagellatefeeding protozoans, which are in turn consumed by copepods, then krill, small fishes, and squids, represent a system that passes on only about a tenth of the energy to apex consumers. Using this theoretical approach, Parsons (1986) described food-web transfers in different environments of the Gulf of Alaska and found that given similar rates of primary production, fjord and shelf waters were capable of supporting twice the apex production of estuarine regions, and approximately 100 times those found in the open ocean. Recent revisions (Welschmeyer et al., 1993) to estimates of

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the open-ocean primary production (adjusted upward by a factor of at least two) still do not bring the deep-ocean apex production close to that on the shelf. Within the seasonal production cycle, food-web efficiency probably change in environments where nutrient limitation shifts the dominant size of primary producers from large to smaller forms in the summer and early fall. When springtime diatom populations are replaced by microflagellates during periods of nutrient limitation, the general trophic efficiency likely declines because a micro-consumer link is inserted into the transfer process. Thus, not only is primary productivity limited by nutrients in most shelf and coastal ecosystems during the summer and early fall, but a more complex food web probably becomes less efficient at transferring energy to higher- level consumers during this same time (Fig. 2.31). There are some exceptions. Larvaceans and pteropods – macrozooplankters adapted to feed directly on very small particles – continue to provide efficient trophic linkages to some apex consumers during the summer and fall. However, this forage base is rarely as abundant or as energy rich as are the large copepods and euphausiids occurring in the spring. To summarize, interaction between upper-layer physics and shelf diatom populations poised to bloom under improving light conditions each year establishes

Figure 2.31: Change in food-chain efficiency in passing energy to seabirds and mammals as dominant primary producers shift from large diatoms during the spring bloom to smaller phytoplankton in the late spring and summer.


the conditions for photosynthetic responses by the plants ranging from short intensive episodes of growth to prolonged periods of lesser and intermittent productivity (Eslinger et al., 2001; Sambrotto et al., 1986). When the upper layer of the ocean stabilizes early and strongly, the resulting pulse of primary production is intense but short-lived because of rapid nutrient diminishment. However, when the upper-layer stability is weak and disturbed by frequent storms, the bloom is much less intense but substantially extended over time. The consequences of these differences are: (1) Under the intense and temporally truncated bloom, the accumulated plant matter is poorly coupled to the pelagic food web, since grazers surviving winter require time to produce their annual broods. Under these “poorly coupled” conditions, a substantial portion of the organic matter sinks out of the surface layers; and (2) With the less intense but prolonged bloom, the pelagic food web is enriched because grazer populations have time to expand and much less organic matter falls to feed seabed consumers. In these quite different ways, organic matter is passed through the food web by the interactions between the living pelagic and benthic assemblages and the physical environment. That growth and reproduction occurs for many consumer stocks in the spring and early summer is evidence of the predictability of the annual diatom bloom

BOX 2.4: HOW DO WE KNOW WHO EATS WHOM? by Theodore Cooney Ever since questions about forage dependencies were first raised, food chains and webs have been constructed mostly by directly describing the stomach contents of consumers. This practice is relatively straightforward but by no means foolproof. For example, digested or partially digested food items are difficult if not impossible to identify. Also, the rate of digestion is not constant for all forage, so what is reported as food by percentage of stomach contents (weight or number) may not be at all correct. In the case of some seabirds, forage dependencies are determined by direct observation – telescopic monitoring of adults provisioning chicks with identifiable food items. More recently, stable isotope analyses have been used to assign “trophic status” using the rule that “you are what you eat,” and specific fatty acid signatures are often traceable to specific prey. These new methods are making it unnecessary to rely entirely on expensive and problematic quantitative stomach analyses. However, none of these techniques by themselves is comprehensive, so marine ecologists seeking information on feeding behaviors and forage dependencies use a suite of different methods.

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BOX 2.5: FEEDING TERMINOLOGY by Theodore Cooney One of the ways that populations of consumers in the sea differentiate themselves is by prey selection. Those that target plankton are termed planktivores, those that select fishes are called piscivores, and those that subsist on a combination of different food types are termed omnivores or generalists. Assigning a species to one of these categories is complicated by the fact that forage dependencies may change seasonally and as an individual moves through its various life stages. Fishes, for example, are mostly planktivores in their larval and juvenile stages, but select other fishes and squids as they mature. Exceptions such as herring and pollock continue to feed on plankton (as adults) when it occurs abundantly (swarms and layers), but also supplement their diets by feeding on fishes when plankton sources are diminished seasonally.

and the short, efficient food webs operating at that time in shelf and coastal environments. But, from midsummer through fall (in all but the open ocean system that is always structurally inefficient), nutrient limitation of primary production and a more complex food-web likely defines less energetically efficient matter transfer – although clearly important to consumers such as the juvenile stages of some fishes – Pacific herring in particular. These seasonal forage shifts suggest planktivory as a primary foraging strategy for consumers early in the season, giving way to omnivory in the summer and fall months, and piscivory during the winter and early spring. In this kind of ecosystem, a generalist such as the walleye pollock is obviously superbly suited to thrive.

2.5. Strategies for Survival 2.5.1. Introduction Alan M. Springer All living things must heed the biological imperative to replace themselves during their reproductive life spans and thus survive as a species. To succeed, plants and animals have evolved life history characteristics, behaviors, and anatomical and physiological adaptations that allow them to find food, conserve energy, keep themselves from becoming the meals of things larger or fiercer, and produce viable offspring.


Strategies for survival vary between species, but all are designed to help individuals take advantage of environmental predictability, cope with environmental uncertainty, and avoid predators. When a strategy fails, it is usually because a species is overcome by forces of a magnitude outside the bounds of common variability and for which it is evolutionarily or circumstantially ill prepared. When this happens, it can place a population or even an entire species at risk. Examples of such strong forces include introduced predators, extreme weather such as El Niño, fisheries, diseases, and contaminants. Many times, causes of change are not conspicuous and are difficult to explain. Because high-level predators in an ecosystem represent an integration and culmination of responses below them, they can be sensitive indicators of ecosystem structure and change. In the following sections, we examine survival strategies of nine prominent species in the Gulf of Alaska, which have experienced major changes in abundance in the past several decades. These focal species – three fishes, three seabirds, and three marine mammals – all exploit the marine environment differently, and all have stories to tell about variability and the nature of change in the Gulf of Alaska. Besides belonging to three different phylogenetic groups, these species span a wide range of physiologies, life history patterns, and trophic levels. Thus, they can serve as sentinels of their environment and its changes. And, in these cases, the reasons their populations have increased or decreased are not entirely certain. By considering their recent histories in the context of their strategies for survival, we may be able to better understand why populations vary and, ultimately, how ecosystems work.

2.5.2. Introduction to Fishes Theodore Cooney In the marine ecosystem, fishes coexist within a matrix of physical constraints and competing biological populations in ways that define their respective niches. These strategies include unique reproductive, growth, and feeding activities that provide competitive leverage for forage resources and ecological space. By exploiting a specific range of physical tolerances and interactions, each species projects a unique “habitat dependency” that describes the conditions necessary for its long-term reproductive success. The life history strategies for marine fishes involve a variety of adaptations, including the life span, age at maturity, fecundity (number of eggs per female), aspects of the early-life history, forage and feeding habits, and reproductive behaviors. The range of adaptation for these features is remarkable. Rockfishes are among the longest living of all fishes – some exceeding 50–70 years of age. Rockfishes are also live bearers (ovoviviparous), but this is the exception among marine fishes. In contrast, a pink salmon is very short-lived – spending but 1 year in freshwater and 1 year in the ocean. The Pacific

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cod (Gadus macrocephalus) is extremely fecund, and relatively long lived, up to 30 years; a female may spawn millions of eggs each year. The eggs of some fishes are pelagic (freely drifting), whereas for others, they are demersal (sticking to rocks or attached plants). Some fishes such as the salmonids and smelts are anadromous – reproducing in freshwater but growing and maturing in the ocean. Some species are planktivores, whereas others are piscivores or omnivores. It is not uncommon for the diets of different life-history stages to differ. Larval and juvenile pollock are principally planktivores, whereas the adults have a more catholic diet, including plankton, fishes, and invertebrates. Some species are even cannibalistic, eating their own young. One fairly common reproductive strategy for fishes is to release large numbers of eggs over a long period each year. Although the probability that any one egg will complete its life cycle to reproduce is vanishingly small, the probability that “some” eggs/larvae/juveniles will find suitable conditions for survival is a certainty. Because the potential for an unusually large survival is present every year for these species, environmental conditions will occasionally allow an “exceptional” number of survivors who will dominate the population for many years. In the case of the Gulf of Alaska, significant variability in year-class strength has been observed in many different fish stocks. In the following, we describe the life history characteristics for three dominant marine fishes – the pink salmon (Oncorhynchus gorbuscha), the Pacific herring (Clupea pallasii), and the pollock (Theragra chalcogramma) – to better understand how each has done in the past, or might do in the future under the various ocean climate regimes in the Gulf of Alaska.

2.5.3. Pink Salmon Theodore Cooney The pink salmon, Oncorhynchus gorbuscha, is the smallest and most numerous of the Pacific salmon. Odd and even year-classes do not interbreed, and with a strict 2-year life cycle, pinks spend about a year maturing in the ocean. Juvenile pink salmon enter the marine environment from freshwater natal areas as juveniles weighing about 0.3 g. After approximately 400 days, they return to spawn and die at an average weight of 1.7 kg; other species of Oncorhynchus grow to much larger adult sizes because of longer periods of feeding and maturation (both marine and freshwater). Over the geographic range of pink salmon in North America, odd and even year-classes trade dominance patterns (Rogers, 1986); runs in Washington and southern British Columbia are typically dominated by odd-year fishes. In southeastern and south-central Alaska, some stocks are dominant in the odd years, some in even years, and others demonstrate a mixture of year-class dominance over time. The very small runs of western Alaska pink salmon are almost exclusively even-year fish. These year-class patterns are unexplained.


In south-central Alaska, the largest runs of pink salmon occur in Prince William Sound, in lower Cook Inlet, and on Kodiak Island. In Prince William Sound, a surprising percentage (>50%) of both odd- and even wild year pink salmon spawn in the inter-tidal reaches of the many small streams and rivers that characterize their natal habitats (Helle, 1970). There is a tendency for odd-year spawners to use the higher regions of these streams and rivers. Each female pink salmon spawns about 1500 eggs before dying (Fig. 2.32). At least two eggs must produce surviving adults to sustain the population of males and females in the absence of a fishery. The large, yolky eggs are deposited in nests dug from the gravels by the females. Spawning usually begins in late June and concludes in September; most occurs in July and August. After a few weeks in the late summer and fall, the eggs hatch and the larval salmon – the alevins – burrow into the sediments to overwinter on yolk-sac reserves. By the following late March and early April, developing juveniles have absorbed most of their yolk and begin emerging, a process that lasts 60 days or more (Bailey, 1969; Taylor, 1988). The emerging juveniles’ transition to saltwater and the year class is entirely in the coastal zone by July (Cooney et al., 1995). There, the juveniles rear for 3–4 months in the shallow edge zone, feeding primarly on zooplankton but also on harpacticoid copepods, insects, polychaetes, and larval fishes (Parker, 1997). In late summer and early fall, surviving juveniles begin moving into the coastal current, initiating a feeding migration that terminates for survivors about 10 months later at their home streams or hatchery. The pink salmon survival strategy places a premium on timing and size. The relatively small number of large eggs per female (compared with many non-salmonid fishes) produce alevins and then juveniles that are much larger than the larvae of pollock and herring. The sedentary period of intertidal and freshwater rearing over the fall and winter months shelters most alevins from predators. However, stream bed scouring from heavy fall rains, freezing, and low oxygen during the winter are all factors that can reduce a year-class during this life stage (Alaska Department of Fish and Game). Pink salmon fry average about 30 mm in total length at entry into marine waters. Surviving juveniles exhibit rates of growth of about 3 to 5% of their body weight per day during early marine residence supported by the gradually warming springtime nearshore ocean and abundant food. A juvenile, entering at 0.3 g, will weigh between 4.5 and 27.0 g after only 90 days in the ocean. Larger size probably allows better escape predators – primarily birds and larger fishes (Willette, 2001). For most fishes, the greatest losses occur in the egg and larval stages. Studies of pink salmon demonstrate this is the case as well. Of the 1500 eggs placed in the redds by a female, some go unfertilized, some are consumed by predatory fishes and birds, some fail to mature correctly, and some of the overwintering alevins succumb to the rigors of the freshwater rearing environment. Heard (1991) estimated that only about 8% of the eggs deposited by female pink salmon develop into viable juveniles. So, hypothetically,

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Figure 2.32: Life cycle of western Prince William Sound pink salmon with alevin rearing areas (1) and migration pathways for out-migrating juveniles (2–4) (top left) and returning adults (5) during the 2-year life cycle. Hypothesized open ocean distributions of maturing adults are shown at the top right, and the sizes and averaged numbers of different life stages surviving from the embryos of a single adult female are depicted in the bottom panel.


of the 1500 eggs deposited in redds by a female, only about 120 alevins survive the winter to enter the nearshore as juveniles, an immediate loss of 1380 potential adults. Perhaps 75% of juvenile pink salmon (wild and hatchery-reared) entering Prince William Sound will be eaten by predators during the first 45 days (Willette et al., 2001). So, after early marine residence and when the juveniles are leaving nearshore waters for the open ocean, just 30 juveniles remain from the original 1500 eggs. The estimated survival to adulthood is about five adults per female spawner in Prince William Sound over the period 1962–1997, according to the Alaska Department of Fish and Game. This level of production has provided for a sustained commercial fishery and generally met desired spawing escapements. The critical survival of juvenile pink salmon in nearshore marine rearing environments depends on food supply and losses to predators (Cooney et al., 2001; Willette et al., 2001). A remarkable concordance between the peak of the juvenile emergence into nearshore waters and the seasonal high in biomass of large calanoid copepods (mostly Neocalanus spp.) suggested that pink salmon in Prince William Sound have evolved a mechanism to enssure plentiful food supply each year for the critical juvenile stage (Cooney et al., 1995). However, further work has demonstrated an even more sophisticated aspect of this timing adaptation. Willette et al. (1999) reported that two of the major predators of juvenile pink salmon – adult walleye pollock and Pacific herring feed almost exclusively on near-surface swarms and layers of large copepods and krill at the same time the young salmon are entering the edge zone. Thus, a kind of predation shelter is established for a few weeks in late April and May that protects the juveniles while they exploit the growth conditions and relative safety of the shallows (Fig. 2.33). In years when the zooplankton layers are not well developed, pollock and herring supplement their diets by feeding more on small fishes, including salmon. Willette (2001) also found that the feeding behavior of juvenile pink salmon occasionally places them at high risk to predation. While there is little evidence that the growth rates of juveniles in Prince William Sound are severely limited by food – marked density-dependent growth is absent except at very high fry abundance – the young salmon apparently prefer the largest copepods available during the spring. When large-bodied Neocalanus or Calanus become scarce in the shallows, the juveniles begin searching over adjacent deeper water and, by doing so, place themselves at greater risk to encounters by adult pollock and herring. In this way, the pink salmon life history strategy for stocks in Prince William Sound protects the juvenile stages early in their marine residence. The apparent strategy is to match the entry timing of the juveniles with the planktonic feeding period of their most important predators – adult pollock and herring. While the larger fishes are feeding on dense surface layers of large calanoid copepods and other macrozooplankters during April, May, and early June, the fry grow in the relatively protected shallow edge of the sound. Water temperature is believed to exert an influence on fry

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Figure 2.33: Relationships between juvenile pink salmon and their food, predators, and amounts of alternative prey for predators – adult pollock and herring. Better juvenile rearing conditions during early marine residence results in higher survivals and better year-class strength (left panels) than when the rearing conditions are poorer (right panels). Red arrows depict predation, with the relative width of the arrow representing importance.

growth at this time. In early June, when the large calanoids begin leaving the upper layers and are no longer available to pollock and herring, the surviving juveniles become alternative prey for other fishes. At this same time, increasing numbers of small gadids (cod and pollock) and adult salmon and Dolly Varden trout also begin invading the warming shallow fry nursery waters. During cool springs, body mass will increase relatively slowly, and the fry will be smaller and more vulnerable in early June. Conversely, when the spring is warmer than average, fry growth rates are expected to be higher. Size-dependent mortality has been demonstrated for pink salmon – the larger juveniles apparently being able to more efficiently avoid predation. In this complex way, temperature-dependent growth and predation sheltering, coupled with the feeding behaviors of the juveniles and their


dominant fish and bird predators determines the survival of a year-class early in the marine life history. The degree to which the early life history strategy exhibited by Prince William Sound pink salmon represents other regions in the northern Gulf of Alaska is unknown. However, since pollock and herring are common and abundant fishes in the coastal zone of the entire Gulf and large calanoids and other macrozooplankters are present in other fjords and sounds, it seems likely that these findings can be considered in a general way until further research proves otherwise. The interested reader is directed to Brodeur et al. (2003) for a thorough review of the latest research on the early life history of Pacific salmon.

2.5.4. Pacific Herring Theodore Cooney Pacific herring, Clupea pallasii, have immense ecological importance and commercial value. These small schooling fish are major contributors to food webs, supporting a wide variety of other fishes, seabirds, and marine mammals in coastal waters (University of Alaska Sea Grant, 2001). Herring are distributed in the eastern North Pacific Ocean from California to the northern Bering Sea (Hay et al., 2001). Separate spawning stocks occur in southeastern, central, and western Alaska. Spawning occurs in the spring, beginning in the south in March and concluding in Norton Sound in June. In Prince William Sound, spawning usually occurs in mid-April at water temperatures close to 4°C. (Fig. 2.34). Each female deposits many thousand small sticky eggs on intertidal and shallow subtidal substrates, including seaweeds and seagrasses. The annual herring spawning draws large numbers of birds, other fishes, and mammals, all targeting the eggs and massed adults in a nearshore feeding frenzy that may last (off and on) for 2–3 weeks. It is not unusual for a spawning population to stretch for tens of miles in the coastal zone, the milt of the males being clearly visible in the water from a distance. For 27 years from 1973 to 1999, the average total distance over which spawning occurred in Prince William Sound was about 90 km (Norcross and Brown, 2001). After spawning, adults feed aggressively to replenish their post-spawned energy reserves. Many move immediately into the surface waters to prey on layers of large calanoids (Neocalanus spp.), euphausiids, and pteropods. Adults supplement their plankton diets when necessary with small fishes, including juvenile salmon (Willette et al., 1999). Unhatched herring embryos are eaten by seabirds, migrating shorebirds, and a variety of fishes and invertebrates exploit the energy-rich egg masses that can also be exposed and further eroded by heavy wave action (Rooper, 1996; Bishop and Green, 2001). As many as 75% of all eggs deposited by spawners can be lost to predation in some years (Rooper et al., 1999).

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Figure 2.34: Life cycle of Pacific herring in Prince William Sound, showing major spawning areas, some proposed larval dispersion paths, and overwintering areas for juveniles and adults (top panel). The sizes and timing of various stages are shown in the bottom panel.


Surviving embryos incubate for approximately 3 weeks in a gradually warming nearshore habitat (Biggs et al., 1992). Hatching begins in mid-to-late May, and the tiny (

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