The Ecology of the Indonesian Seas: Part 1 [7, 1 ed.] 9625930787, 9789625930787

Table of contents :
Intro
Table of Contents
EMDI
Acknowledgements
Chapter 1: Introduction
Chapter 2: Geology
Chapter 3: Physical Oceanography and Metereology
Chapter 4: Introducing Coral Reefs
Chapter 5: Geological History of Reefs
Chapter 6: Coral Reef Origins: The Theories
Chapter 7: Scleractina: The Reef-Builders
Colour Plates
Chapter 8: Non-Scleractinian Cnidaria
Chapter 9: Foraminiferida
Chapter 10: Environmental Factors
Chapter 11: Coral Reefs: Natural Disturbances
Chapter 12: Coral Reefs: Growth and Development

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T H E ECOLOGY OF T H E I N D O N E S I A N SEAS PART I CHAPTERS

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T H E ECOLOGY OF INDONESIA SERIES VOLUME VII

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T H E ECOLOGY OF I N D O N E S I A S E R I E S

Volume VII: The Ecology of the Indonesian Seas, Part One Volume VIII: The Ecology of the Indonesian Seas, Part Two Other titles in the Series Volume I: The Ecology of Sumatra Volume II: The Ecology ofJava and Bali Volume III: The Ecology of Kalimantan Volume IV: The Ecology of Sulawesi Volume V: The Ecology of Nusa Tenggara and Maluku Volume VI: The Ecology of Irian Jaya Produced by Environmental Management Development in Indonesia Project, a cooperative project of the Indonesian Ministry of the Environment and Dalhousie University, Halifax, Nova Scotia under the sponsorship of the Canadian International Development Agency This book is dedicated to our parents.

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Indonesia

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The Ecology of the Indonesian Seas PART I CHAPTERS

1-12

TOMAS TOMASCIK ANMARIE JANICE MAH ANUGERAH NONTJI MOHAMMAD KASIM MOOSA

PERIPLUS EDITIONS

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Copyright © 1997 Dalhousie University All rights reserved Published by Periplus Editions (HK) Ltd. ISBN 962-593-078-7 , Publisher: Eric Oey Typesetting and graphics: JWD Communications Ltd. Distributors: Australia: University of New South Wales Press Ltd Sydney NSW 2052 Indonesia: C.V.Java Books Jalan Kelapa Gading Kirana, Blok A14 No. 17 Jakarta 14240

Japan:

Tuttle Shokai Ltd 21-13, Seki 1-Chome, Tama-ku, Kawasaki, Kanagawa 214 Singapore and Malaysia: Berkeley Books Private Ltd. 5 Little Road #08-01 Singapore 536983 United Kingdom: Oxford University Press Great Clarendon Street Oxford, OX2 6DP United States: Charles E. Tuttle Co., Inc., RRI Box 231-5, North Clarendon, VT 05759-9700

Printed in the Republic of Singapore

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Table of Contents (Parti)

EMDI xi Acknowledgements xiii Chapter 1 Introduction 1 Where Are the Indonesian Seas? 1 Historical Perspective 4 Objectives of the Book 6 Main Themes 6 Chapter 2 Geology 9 Introduction 9 Structure of the Earth 9 Mantle Convection 11 Plate Tectonics 12 Earthquakes 21 Volcanoes 25 The Shapes of Volcanoes 25 Other Volcanic Landforms 26 Factors Determining the Types of Volcanic Activity 28 Products of Volcanic Activity 29 Distribution of Volcanoes 31 Distribution of Active Volcanism in Indonesia 32 Indonesia's Most Famous Eruptions 33 Tectonics of the Indonesian Archipelago 36 Introduction 37 Plate Motions 37 Historical Setting 38 The Early History of the Indonesian Archipelago 39 T h e S u n d a A r c 41 TheBandaArc 42 The Banda Arc and Plate Motions 43 Sulawesi 44 The Wallace Line 46 Geological History and Biogeography 48 Basins 49

O c e a n Basins - F o r m a t i o n to Demise 49 T h e F o r m a t i o n of t h e C u r r e n t Ocean Basins 50 Marginal Seas 54 Sedimentary Basins 56 Classification 56 Paleogeography of the Indonesian Archipelago 62 Late Jurassic - Cretaceous (160-65 Ma) 62 Paleocene - Eocene (65-36 Ma) 62 Oligocene (36-25 Ma) 62 Early to Middle Miocene (25-12 Ma) 63 Late Miocene to Recent (12 Ma to Present) 63 Postscript 64 Chapter 3 Physical Oceanography and Meteorology 71 Introduction 71 Significance of Indonesian Seas 71 Ocean Circulation 72 Physical Oceanography 72 Geography and Bathymetry 72 Meteorology 94 I n f l u e n c e s of G e o g r a p h y a n d Geology 94 The General Climate 94 The Monsoon Seasons 98 Trade Winds 99 Variability 99 Ecological Significance of Physical Processes 100 Ecological and Physical Processes 100 Physical Oceanography 100 Meteorology 101 Chapter 4 Introducing Coral Reefs 109 Introduction 109 Coral Reefs under Threat 112

Distribution of Coral Reefs 115 World Distribution Patterns 115 Coral Reefs of the Atlantic Ocean 118 Coral Reefs of the Indian Ocean 119 Coral Reefs of the Pacific Ocean 120 Regional Perspective 123 Coral Reefs of the Indo-Pacific Region 123 The Nature of Coral Reefs 124 Coral Reef Definitions 126 Coral Reef Definitions for Management 132 Global Coral Reef Area Estimate 134 Areal Estimates of Coral Reefs in Indonesia 136 Chapter 5 Geological History of Reefs 145 Introduction 145 Fossil Reefs 147 The Early Years 149 The Precambrian Reefs 149 Stromatolites 149 Extant Stromatolites in Indonesia 151 The Paleozoic Reefs 166 Archaeocyatha 166 Porifera and Stromatoporoidea 167 Tabulate Corals 169 Rugose Corals 171 Corals and Stromatoporoidea 171 Paleozoic Fossil Record 172 The Permian Extinction 177 The Mesozoic Reefs 178 Triassic 181 Jurassic 183 Cretaceous 185 Cretaceous Extinction 189 Cenozoic Reefs 189 The Paleogene 190 The Neogene 192 The Quaternary 197 Chapter 6 Coral Reef Origins: The Theories 207 Introduction 207

viii

Theories of Coral Reef Origins 210 Darwin's Theory of Subsidence 211 Challenges to Darwin's Theory 216 Chapter 7 Scleractinia: The ReefBuilders 233 Introduction 233 The Anatomy of a Reef-Builder 233 Coral Animal: The Polyp 235 Reproduction 246 Coral-Zooxanthellae Symbiosis 251 The Animal Host 252 Calcification and CaCOg Deposition 253 Hermatypic or Zooxanthellate? 255 Classification of Scleractinia 256 Class Anthozoa 256 Subclass Zoantharia 257 Order Scleractinia 257 Morphological Characteristics 260 Classification of Scleractinia 267 Major Subdivisions 267 Family Astrocoeniidae 268 Family Pocilloporidae 268 Family Acroporidae 269 Family Poritidae 279 Family Siderastreidae 280 Family Agariciidae 280 Family Pectiniidae 281 Family Oculinidae 282 Family Fungiidae 283 Family Mussidae 288 Family Faviidae 289 Family Trachyphylliidae 293 Family Merulinidae 293 Family Caryophylliidae 293 Family Dendrophyllidae 294 Coral Diversity Patterns 296 Deep-Water Scleractinians 316 Nutrient Pathways 321 Tracing Food Webs with Stable Isotopes 321 Coral Reef Productivity 322 Chapter 8 Non-Scleractinian Cnidaria 325 Introduction 325 Classification of Subclass Octocorallia 327

Chapter 10 Environmental Factors 421 Introduction 421 Temperature 421 Temperature and Reefs 421 Temperature and Corals 423 Coral Reefs and Upwelling 426 High Temperatures 429 Salinity 429 Definitions 429 Salinity and Global Reef Definition 432 Salinity and Indonesian Reef Distribution 435 Effects of Low Salinities 439 Berau Islands: Case Study 443 Effects of High Salinities 446 Light and Coral Reefs 447 Light Environment in Indonesia 450 Sedimentation 468 Definitions 468 Sediment Transport, Origin and Effects 469 River Transport 469 Sediment Origins 470 Ambon Bay: Case Study 471 Tanjung Setan: Case Study 473 Effects of Large Rivers 474 Land Erosion 476 Sedimentation, Corals and Coral Reefs 477 Sediment Rejection by Corals 478 Coral Recruitment 480 Suitable Substrates 480 Nutrients 481 Eutrophication 483 Circulation 484

Reef-Building Contribution of Octocorallia 332 Octocoral-Zooxanthellae Symbiosis 334 Class Hydrozoa 341 Introduction 341 Order Milleporina 347 Order Stylasterina 348 The Non-Skeletal Anthozoa 350 Introduction 350 Chapter 9 Foraminiferida 371 Introduction 371 Classification 374 Biological Aspects 374 Test Morphology 374 Reproduction 378 Symbiosis 380 Calcification 384 Heterotrophic-Feeding 385 Benthic Foraminifera 388 General Distribution 388 Economic Value of Foraminifera 389 Bali Case Study 389 Source of Sanur Beach Sand 395 Foraminifera as Environmental Indicators 397 Species Associations 397 Subpolar Forams in Indonesia 399 Deep Chlorophyll Maximum Layer 401 Morphological Indicators 402 Test Porosity 404 Test Coiling 404 The Use of Stable Isotopes 405 Stable Oxygen Isotopes 405 Ice Volume Effect 407 Measurement 408 Use of 5 18 0 in Ontogenic Studies 409 Stable Carbon Isotopes 411 Paleoceanography and Paleoclimates 412 Northern Molucca Sea 413 Makassar Strait 414 BandaSea 415 Contribution to Marine Sediments 416

Chapter 11 Coral Reefs: Natural Disturbances 487 Introduction 487 Atmospheric Disturbances 487 Monsoonal Storms 488 Tropical Cyclones 493 Cyclone "Lena" (January 23, 1993) 495 El Nino-Southern Oscillation (ENSO) 499

IX

Rising Sea Levels 608 Stable Sea Level 612 Falling Sea Levels 614 Destructive Processes 614 Physical Processes 615 Biological Processes 618 Reef Microborers 621 Boring Sponges 623 Boring Worms 624 Boring Molluscs 630 Echinoids 635 Herbivorous Reef Fish 637

El Nino and Coral Reef Bleaching 500 Reef Bleaching Global Patterns 501 Reef Bleaching: Trigger Mechanisms 502 Zooxanthellae Expulsion 503 Nature of Coral Bleaching 504 UV Radiation 507 Coral Bleaching and Sea Level 510 Physiological Impacts of Bleaching 516 Coral Community Response 521 Volcanism, Earthquakes and Tsunamis 522 Volcanism 522 Geographic Distribution of Volcanoes 523 Volcanism and Coral Reefs 524 BandaApi: Case Study 529 Banua Wuhu (Mahengetang) Underwater Volcano 551 Shallow-Water Hydrothermal Vents 558 Earthquakes 565 Coral Reefs and Earthquakes 567 Coral Reefs in Sunda Strait 567 E a r t h q u a k e s a n d B a n d a Api Eruption 568 Maumere Bay Earthquake 571 Tsunamis 572 Generating Forces 572 Physical Properties of Tsunamis 573 Tsunamis in Indonesia 573 The Krakatau Tsunami 574 The Maumere Bay Tsunami 574 The Banyuwangi Tsunami 576 Biological Perturbations 577 Chapter 12 Coral Reefs: Growth and Development 579 Introduction 579 Framework for Classification 580 Shelf Reefs 580 Classification of Shelf Reefs 586 Fringing Reefs 589 Oceanic Reefs 594 Sea-Level Fluctuations 598 Reef Growth and Development 605 Reef Accretion Rates 606 Reef Development 607 x

EMDI The Environmental Management Development in Indonesia Project (EMDI) was designed to upgrade environmental management capabilities through institutional strengthening and human resource development. A joint project of the Ministry of State for Environment (LH), Jakarta, and the School for Resource and Environmental Studies, Dalhousie University, Halifax, Nova Scotia, EMDI supported LH's mandate to provide guidance and leadership to Indonesian agencies and organizations responsible for implementing environmental management and sustainable development. Linkages between Indonesian and Canadian organizations and individuals in the area of environmental management were also fostered. EMDI received generous funding from the Canadian International Development Agency (CIDA). CIDA provided Cdn$2.5 million to EMDI-1 (1983-86), Cdn$7.7 million to EMDI-2 (1986-89), and contributed Cdn$37.3 million to EMDI-3 (1989-95). Significant contributions, direct and in kind, were made by LH and Dalhousie University. EMDI-3 emphasized spatial planning and regional environmental management, environmental impact assessment, environmental standards, hazardous and toxic substance management, marine and coastal environmental management, environmental information systems, and environmental law. The opportunity for further studies was offered through fellowships and internships for qualified individuals. The books in the Ecology of Indonesia series form a major part of the publications programme. Linkages with NGOs and the private sector were encouraged. EMDI supported the University Consortium on the Environment comprising Gadjah Mada University, the University of Indonesia, the Bandung Institute of Technology, the University of Waterloo, and York University. Included in EMDI activities at Dalhousie University were research fellowships and exchanges for senior professionals in Indonesia and Canada, and assistance for Dalhousie graduate students undertaking thesis research in Indonesia. For further information about the EMDI project, please contact: Director School for Resource and Environmental Studies Dalhousie University 1312 Robie Street Halifax, Nova Scotia Canada B3H 3E2 Tel. 1-902-494-3632

XI

xii

Acknowledgements In a book of this dimension it is difficult to individually thank each person, and we hope that we are forgiven for any omissions. Our thanks are extended to all those in the Ministry of Environment, Indonesia, and the Canadian International Development Agency who made this book project possible. Special thanks to all the advisors, management, and staff of the EMDI Project. We would especially like to thank Shirley Conover, Aca Sugandhy, Hadi Alikodra, Sudaryono, George Greene, Gerry Glazier, Ray Cote, John Patterson, Clifton Potter, Pauline Lawrence, Brian Yates, Diane Blanchford, Barbara Patton, Susan Woods, Gerard Belanger and Antin Wibowo. Our profound thanks to all at Dalhousie University and at the School for Resource and Environmental Studies who have assisted us with literature surveys and acquisitions. Our deep gratitude to all our colleagues who made a significant contribution to this book. Their contributions are clearly acknowledged in the text of the book. We would like to express our thanks to Alan Logan, Lyndon Devantier, Howard Spero, Peter Bell, Rob Van Woesik, John Boers, Koos J.C. den Hartog and Leen van Ofwegen for their reviews of the manuscript. Special thanks to John Clark for many suggestions and encouragements. Identification of the species in this book, many of them new discoveries, often necessitated the expertise of taxonomists. For their assistance we wish to thank G. Allen, R. van Soest, P. Ng, CI. Massin, J. van der Land, J.C. den Hartog, L. van Ofwegen, C. Fransen, J. Goud, C. Kishinami, B. Collette, M. Hoogmoed, P. Cornelius, A. Allison, J. Goud, E. Verheij, C. Bryce, AJ. Bruce, N.L. Bruce, L. Hillis, L. Marsh, P. Mather, W. Ivanstoff, H. Larson, O.H. Arinardi and Trimaningsih. Our deep and special gratitude goes to Roger Steene, Ron and Valerie Taylor, Chuck Birkeland, Gerald Allen, Koos den Hartog, Charles Fransen, Leen van Ofwegen, The National Museum of Natural History in Leiden, Robert van Soest, Mark A. Johnson, Mr. Fernandez, Barbara Brown, William Patzert and SPOT Image Corporation, for providing us with wonderful photographs that have brought this book to life. We offer sincere apologies to those whose beautiful colour slides had to be converted to black-and-white prints. The book also greatly benefits from the wonderful artwork of Basuki Rahmad and Mickey Meyer. We would like to express our gratitude to Mr. Dicky Daryanto and Garuda Indonesia, as well as Mr. Permadi and Aneka Herman of Sempati Air for travel assistance. Our special thanks to the following individuals and groups which have provided support for our research activities: Ron Holland, Graham and xiii

Donna Taylor of Borneo Divers; Paul Sugiono and Peter Sugiono of P.T. Sangalaki Resort; Des Alwi, Tania Alwi and Al Welsh in the Banda Islands; Quark Expeditions, Banda Sea; Mr. Fernandez of Sao Wisata, Flores; Henny Batuna of Murex in Manado; Ibu Sita Wachjo of the Yaysan Sumber Daya Laut, Lembata Island; and Kal Muller in Flores. The completion of the book would not have been possible without the help of many friends who offered encouragement and suggestions to improve the quality of this book and made our stay in Indonesia a memorable experience. Our deep gratitude goes to Al and Anita Welsh, Keith James, Debra Nishida, Suzanne Gendron, Mickey Meyer, Tokkie and Jacqueline Elliot, Sandra van Woesik, Wawan Kiswara, Dwi Sasongko, Errol Billing, Tania Alwi, Rili Djohani, Gerry and Marlene Glazier, John and Sheila Patterson, Robin Harger, Ibu Sita Wachjo, and Kent Wiley. Finally we would like to express our deep gratitude to everyone involved in this project.

xiv

Chapter One

Introduction Situated upon the equator, and bathed by the tepid water of the great tropical oceans, this region enjoys a climate more uniformly hot and moist than almost any other part of the globe, and teams with natural productions which are elsewhere

unknown. — WALLACE 1869

WHERE ARE THE INDONESIAN SEAS? The tropical Indo-Pacific ocean has recently been considered as the 'largest ecological system on earth' extending from the eastern tropical Pacific to the east coast of Africa (Sheppard et al. 1992). This vast area contains much of the world's marine biodiversity including a wide range of marine and coastal environments that support it. The Southeast Asian waters, and those of New Guinea, biogeographically known as the Indo-Malayan triangle, are located in the middle of the Indo-Pacific, and represent an area of high marine biodiversity. Ekman (1953) considered this area as the faunistic centre from which other regions of the Indo-West Pacific recruited their faunas. Briggs (1974) suggested that the Indo-Malayan triangle with its widely distributed marine fauna covers a broader area within the IndoPacific, known as the Indo-Polynesian Province, with the Indo-Malayan triangle as the centre. Indonesia, derived from the Greek words "Indos" meaning Indian and "nesos" meaning islands, is the main part of the Indo-Malayan triangle, stretching roughly from 6° N to 10° S and from 95° E to 142° E. Table 1.1 provides a brief geographic summary of the Indonesian Archipelago. Indonesia occupies a central position of the Indo-Pacific, thus creating a permeable barrier between the Pacific and Indian Oceans and the Asian and Australian continents. With more than 17,500 islands and a coastline in excess of 80,000 km, the Indonesian Archipelago is a storehouse of marine biodiversity. The marine and coastal environment of Indonesia includes a high diversity of ecosystems such as beaches, sand dunes, estuaries, mangroves, coral reefs, seagrass beds, coastal mudflats, tidal forests, algal beds as well as many small island ecosystems. In addition to these critical coastal ecosystems, the Indonesian Archipelago contains vast continental slope areas, abyssal plains and deep oceanic trenches. Each of these marine ecosystems, with its associated habitats, supports a wealth of marine resources. About 78% of Indonesian territory are shallow seas located on the Sunda and Sahul Shelves, separated by the deep Timor, Banda and Flores Seas (fig. 1.1). Most of the present, albeit limited, 1

2

INTRODUCTION

knowledge of marine life comes from shallow water biota at depths of less than 200 m. As the world's largest archipelago, Indonesia is positioned in a strategic location with respect to global ocean circulation patterns. The dynamic nature of the archipelagic seas, their interaction with the Pacific and Indian Oceans and the monsoonal climate to a great extent explain the high marine biodiversity of the region. Plate tectonics, and the associated seismicity and volcanism, played, and continue to play, a key role in the geologic evolution of the archipelago. These natural geological processes are the most important long-term factors affecting the physical, chemical and biological processes as well as human populations inhabiting the archipelagic islands as early as the Paleogene. The movement of the crustal plates is a dynamic process that has important ramifications for the present-day geographic distribution of marine ecosystems and associated biological communities. Since the climate of the archipelago is under the influence of the Asian-Australian monsoon system, the climate exerts a major influence on the large-scale circulation patterns of the intra-archipelagic seas and plays a significant role in the productivity of the coastal and marine systems.

Table 1.1. The geographic summary of the Indonesian Archipelago. Parameter

Unit of measurement

Total number of islands

17,508

Coastline length (baseline)

80,791 km

Total land area Area of archipelagic (inner) seas

1,926,337 km2

2,820,000 km2 420,000 km

2

Continental Shelf area

1,500,000 km

2

Area of EEZ (Exclusive Economic Zone)

2,730,000 km

2

Area of territorial (12-nm zone) seas

Total area of national jurisdiction

Notes Major islands: Sumatra, Java, Sulawesi. Major segments of Borneo (Kalimantan), and New Guinea (Irian Jaya)

The actual length of the Indonesian coastline may be about 204,000 km (Astuti et al. 1994)

24.4% of total area under Indonesian jurisdiction 35.7% of total area under Indonesian jurisdiction 5.3% of total area under Indonesian jurisdiction 19% of total area under Indonesian jurisdiction 34.6% of total area under Indonesian jurisdiction

7,892,350 km2

Source: Soegiarto and Polunin 1981; Astuti et al. 1995.

Figure 1.1. Map of the Indonesian Archipelago showing the extent of Sunda and Sahul Shelves as delineated by the 200-m isobath.

WHERE ARE T H E INDONESIAN SEAS? 3

4

INTRODUCTION

HISTORICAL PERSPECTIVE Studies of Indonesian marine fauna and flora pre-date Linnaeus' Systema Naturae (1758 - 10th edition). The well-known naturalist, G. E. Rumphius (1627-1702), worked on materials collected from Ambon, and other areas of the Moluccas, with some materials likely originating from other parts of the eastern Indonesian seas. Working with the VOC and stationed in Ambon, Moluccas, most of his life, his contribution to science was monumental. His work was collated in two great works, D'Amboinsche Rariteitkamer (1705) and Herbarium Amboinense (1741-50); the latter work catalogued 1200 species of marine plants (fig. 1.2). The rich marine diversity of the Indonesian Archipelago has motivated marine scientists to organize numerous expeditions to this fascinating region. Early French expeditions include Physicienne (1817-20); Coquille (1822-25), Astrolabe (1826-29) and Bonite (1836-37). English expeditions during this period included the Beagle (183236) with Charles Darwin, Sulphur (1836-42), and Samarang (1843-46). These were followed by the Challenger Expedition (1873-76), which passed through the eastern part of the archipelago and concentrated on oceanography as did the German expedition Valdivia (1898-99). Challenger's 50 volumes, written over a period of 20 years, made a great contribution to our understanding of the world's oceans. One of the most significant early expeditions to the eastern parts of the Indonesian Archipelago was that of the Siboga (1899-1900). These early expeditions generated great interest in the "Malay Archipelago," not only in the marine environment, but also in terrestrial exploration. Indeed, the period between 1850 and 1900 saw an influx of terrestrial naturalists, collectors and travelers. The works of the great naturalist Sir Alfred R. Wallace (1854-62) contributed greatly to our understanding of the biogeography of the archipelago, and his faunal and floral demarcation known as "Wallace's Line" was later explained by plate tectonics. In 1842, Pieter Bleeker, a prominent ichthyologist, began working on Atlas Ichthyologique which constituted the first ichthyological research in Indonesia. It was not until 1904, with the establishment of the Visscherij Station in Pasar Ikan (Jakarta), that marine biology became entrenched in the national development planning. In 1919, the Visscherij Station was given a new mandate for much broader oceanographic research, and was renamed Laboratorium voor het Onderzoek der Zee.

The 1900s began with the German Expedition Planet (1906-07), followed by the highly successful Snellius Expedition (1921-30). However, data from these early expeditions have not been widely available. More recent expeditions were aboard the Russian research vessel Vityaz (1963), the American Vega (1963-64), the Baruna Expedition (1964), the Mariel King Memorial Expedition limited to the Moluccas (1970), the Rumphius Expeditions I-IV (1971, 1973, 1977 and 1980), the FrancoIndonesian Corindon cruises (1981,1983), the Snellius II Expedition (1984-85) and the Franco-Indonesian Karubar Expedition (1991). Considering the huge amounts of information collected by the various expeditions, it is surprising that the marine biodiversity of the Indonesian Archipelago remains largely unknown and generally unavailable to the public. The results of many expeditions are frequently reported as parts of other works, or they are published several decades following the expedition. For example, it was not until 1982

HISTORICAL PERSPECTIVE

5

Figure 1.2. One of the first illustrations of a marine alga from eastern Indonesia shown in Herbarium Amboinense (1750). Drawing by Mickey Meyer.

that information on the majid crabs collected during the 1929-30 Siboga Expedition and the 1970 Mariel King Memorial Expedition in the Moluccas was published (Griffin and Tranter 1982). Forest (1987) cited specimens of pagurid crabs collected in the Makassar Strait during the Corindon II Expedition in his discussion of the Polychelidae from other parts of the world. Specimens from these Indonesian cruises made valuable contributions to world biogeography and taxonomy. Today the main centre for oceanographic research is at P 3 0 (Pusat Penelitian dan Pengembangan Oseanobgi) located in Ancol, Jakarta, which is the successor of the Lembaga Oseanologi Nasional (LON), with a new mandate that focuses on basic oceanographic research as well as on applied research in all fields of marine sciences, from engineering to taxonomy.

6

INTRODUCTION

OBJECTIVES OF THE BOOK The main objective of The Ecology of the Indonesian Seas is to introduce Indonesian students to the fascinating marine environment of the Indonesian Archipelago. The book has been written as a teaching text suitable for undergraduate courses, government agencies and the interested public, and is an introduction to the marine ecology of the Indonesian Archipelago, one of the world's most dynamic marine and coastal regions. This book provides an overview of a number of coastal and marine environments of the Indonesian Archipelago, with a focus on ecological processes. However, the book primarily focuses on shallow-water coastal ecosystems, particularly coral reefs. The influence of human populations on the marine and coastal resources and their supporting ecosystems are also examined.

MAIN THEMES The overall theme of the book is to touch upon our current knowledge, of the ecology of the archipelagic seas in view of its geologic history, the physico-chemical and biological environments, and human interactions. Marine ecology is viewed as a study of the interrelations between living organisms, including human populations, and their environment. The chapter on the physical oceanography and climate of the region illustrates the geographic complexity of the archipelago, and the key role of the monsoons in the overall physical characteristics of the marine environment. Major global climate changes associated with glaciation had a pronounced influence on the distribution of terrestrial biota; however, this influence is less clearly exhibited in the marine realm, where few barriers to dispersion exist. The book places great emphasis on the geologic history of the archipelago, since the diversity of marine ecosystems and associated biological diversity are a function of geologic history. Specifically, plate tectonics, seismicity and volcanism of the region directly influence the physical distribution of landmasses and submarine topography, which in turn influence the overall circulation patterns of the archipelagic seas and the climate itself. Recent advances in physical oceanography, through cooperative research studies with international scientific organizations, have highlighted the key role of the Indonesian Throughflow, a major transport of Pacific water masses into the Indian Ocean through the archipelagic basins. The dynamic nature of the archipelago, in terms of its geology and climate, directly influences the inputs of nutrients and trace elements into the seas through volcanic activity, sedimentation, and in recent history, large amounts of soil eroded from areas under human influence. The sections on biological communities collate much information on the various marine communities, but more importantly they point to serious gaps in our knowledge. While the need for applied marine research is clearly recognized, this book points out that a serious lack of basic research is now making advances in applied fields very difficult. A more balanced approach is required if the government's sustainable-use policies of marine and coastal resources and the conservation of biodiversity are to be successfully implemented.

MAIN THEMES

7

Marine resources, both flora and fauna, have been exploited for human consumption or other needs, since the early history of mankind. Fish, crustaceans, molluscs and marine mammals are appreciated and widely consumed as a source of animal protein, or as a delicacy, by many cultures throughout the world. Coral reef fish and other reef inhabitants displayed in the world's public aquariums attract millions of people, and are also enjoyed by people in the privacy of their own homes. In Indonesia, marine resources are a valuable export commodity. Corals, shells, and seaweeds are exploited as industrial raw materials. People are now looking to the sea for bioactive substances for medical and pharmaceutical purposes. Scientists know very little about the loss of genetic and species resources from marine environments. Marine species, or even populations, have disappeared in historical times. Major disturbances to marine ecosystems could lead to the loss of genetic diversity. Marine, and especially coastal areas, are affected by human activities onshore and inland. A solid knowledge and understanding of the marine and coastal environments of the Indonesian Archipelago must form the foundation for effective management and conservation of the marine and coastal resources. It is the objective of this book to provide the initial tools from which Indonesia's most precious asset, its people, may manage and conserve for future generations, its second-most valuable asset, the Indonesian seas.

8

Chapter Two

Geology Geology teaches us that the surface of the land and the distribution of land and water is everywhere slowly changing. It further teaches us that the forms of life which inhabit that surface have, during every period of which we possess any record, been also slowly changing.—WALLACE 1869

INTRODUCTION The world changes, but to most of us, our Earth is the symbol of stability. Viewed from the standpoint of a human life span, and indeed generations of humans, the foundations of Earth remain unchanged. However, on a geological time scale, the earth has undergone radical transformations from its birth to present. These slow but ongoing changes are modifying the surface of the earth, reshaping the continents and ocean basins. More rapid changes are occurring on the earth's surface as a result of human activity. Examples of anthropogenic actions which have caused major modifications to the earth's topography include the construction of dams, open-pit mining, and rice cultivation. Flannery (1994) suggests that the 60,000 years of human occupation in Australia have impacted the Australian continent to such an extent that virtually all of the continent's ecosystems are in some sense manmade. In Indonesia, Javanese and Balinese rice cultivation have had a major impact on the ecology of those islands (see Whitten et al. 1996), as had open-pit mining on the islands of Sulawesi and Irian Jaya. But human impacts on the Indonesian Archipelago are the topic of a later chapter. Of interest to us now is the land beneath our feet, its birth and evolution. Several theories evolved during the 1900s leading to the now accepted theory of plate tectonics. However, prior to discussing plate tectonics and its role in shaping the Indonesian Archipelago, we will briefly review the structure of the earth.

S T R U C T U R E OF T H E E A R T H The earth has a layered structure comprising the core, mantle, and crust (fig. 2.1). The core is divided into a solid inner core and a nickel-iron outer core. The outer core is liquid, and it is the circulation within the outer core that generates the earth's magnetic field. This point will be significant when we later discuss magnetic anomalies. 9

10

GEOLOGY

Figure 2.1. The earth's layered structure. Cross-section through one hemisphere of the earth showing the inner core (1), the outer core (2), the lower mantle (3), the upper mantle (4), and the crust (5). The inner core is approximately 2414 km in diameter. The outer core is 2253 km thick, as is the lower mantle, the upper mantle is roughly 644 km in thickness, and the crust varies from 3.2 km under parts of the oceans to 120 km thick beneath mountains. Figure not to scale.

The mantle is mostly solid, consisting of dark, heavy rocks which are rich in iron-magnesium silicates such as olivine and pyroxene (Rhodes 1991). The asthenosphere is part of the upper mantle. It is known as the zone of mobility, since it is "nearly molten" (Gross 1990) with convection movements and isostatic adjustments occurring within this layer. Magmas (molten rock) may also be generated (Bates and Jackson 1980). The lithosphere includes the crust and part of the upper mantie. It is 50-100 km thick. The lithosphere floats on the asthenosphere. The crust, or what we know as the earth's "surface", forms die upper portion of the lithosphere. There are two types of crust: continental and oceanic. Continental crust is generally 30-40 km thick, but may be 120 km thick beneath mountains and is rich in aluminum and silica. Oceanic crust is 3-7 km thick and rich in magnesium and iron. Continental crust is thicker but less dense than oceanic crust, so it floats higher in the asthenosphere than oceanic crust. This is an example of isostasy, an equilibrium condition comparable to floating. However, there are two different concepts of the mechanism of isostasy: the Airy hypothesis and the Pratt hypothesis (fig. 2.2). The Airy hypothesis postulates an equilibrium of crustal blocks of the same density but of different thickness (i.e., topographically higher mountains are of the same density as other crustal blocks, but have greater mass and deeper roots) (Bates and Jackson 1980). This is an example of floatational equilibrium (Spencer

11

Increasing density

STRUCTURE OF THE EARTH

Figure 2.2. Isostasy: The Airy hypothesis and the Pratt hypothesis. The Airy hypothesis (A) is based on the premise that the crustal blocks are of equal density but the roots are at different levels. The Pratt hypothesis (B) assumes that the blocks are of different densities but the differences are compensated at a certain depth. The Heiskanen hypothesis (C) is a combination of the Airy and Pratt hypotheses. Spencer 1972, Kingston

1988.

1972) and can be compared to blocks of wood floating in water. The Pratt hypothesis postulates an equilibrium of crustal blocks of different densities (i.e., topographically higher mountains are less dense than topographically lower units such as the ocean floor, while the depth of crustal material is everywhere the same) (Bates and Jackson 1980). The depth at which the effects of the different densities are balanced (i.e., the level to which all the blocks sink) is the level of compensation. The lesser-known Heiskanen hypothesis combines assumptions from Airy and Pratt to account for mountain roots, and the variations in crustal densities. It is currently thought that neither the Airy nor the Pratt hypothesis can fully explain the isostatic equilibrium which exists. The lithosphere is broken into several pieces or "plates" and it is the movement of these plates which is known as plate tectonics. Mantle Convection Mantle convection is responsible for the movement of the lithospheric plates. Since the earth's crust is a part of the mantle, it is carried "piggyback", with the mantle's movements. The mantle is warmed at the core-mantle boundary, and then slowly rises to the surface where it cools and sinks again. Cooling of the mantle takes place through several processes: • volcanic eruptions at mid-ocean ridges and hot spots; • seawater circulation through new crust; • heat conduction through the ocean floor (Gross 1990).

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Once cooled, the thickened and denser lithospheric plates are eventually drawn down through subduction trenches, coming to rest on the core-mantle boundary. Over a period of millions of years, the plates will warm and decrease in density until they are sufficiently buoyant to rise through the mantle and begin the cycle again. Images of rising and sinking mantle can be seen using seismic tomography which is similar to computerized axial tomography or the "CAT scans" used in the field of medicine to image human organs (Gross 1990). Hot Spots. Some spots on the core-mande boundary are anomalously hot. At these hot spots, plumes of molten rock rise through the mantle and form volcanoes on the surface. Hot spots do not move with the overlying mantle and lithospheric plates. They remain in the same location for tens of millions of years, recording the movement of the overlying plate with a chain of volcanoes (Gross 1990). The Hawaiian islands and the Emperor Seamounts were formed over a hot spot which is currendy found to the southeast of the Hawaiian islands. Loihi is still submerged, but will eventually join the chain of volcanic islands. There are no hot spots in the Indonesian Archipelago. Hot spots are also believed to be responsible for the formation of several aseismic ridges such as the Ninety-east Ridge in the Indian Ocean, and the Walvis and Rio Grande Rises in the South Atlantic (Brown et al. 1989). Plate Tectonics The Beginnings of the Theory of Plate Tectonics. The theory of continental drift (or more appropriately, continental displacement) was first proposed in 1915 by Alfred Wegner after he observed the jigsaw puzzle fit of the eastern coast of South America with the west coast of Africa. However, he was not the first to observe this intriguing fact, for Alexander von Humbolt, the naturalist-explorer, had made a similar observation in 1801 (Seibold and Berger 1982). Wegner's book, The Origin of Continents and Oceans, was originally published in 1915 in German, followed by the English translation in 1924. Wegner contended that the similarity of the rocks and fossils of the two coasts supported his theory. He theorized that all the continents were once part of a single continent known as Pangaea. The continents moved apart by the displacement of large plates of continental (sialic) crust, moving freely across a substratum of oceanic (simatic) crust (Bates and Jackson 1980). However his hypothesis was not widely accepted by the scientific community. The major objections were due to his proposed mechanisms. Wegner suggested that the continents were rigid plates moving through the ocean basins. The motions were driven by the variations in the gravitational attraction of the earth's equatorial bulge and the westward drift due to the attractions of the Sun and the Moon (Gross 1990). In addition, Wegner's fossil evidence was not definitive. Prior to the first mapping of the ocean floor in the 1950s, it was believed that the ocean basins and continents were stable features of the earth. Robert Dietz, in 1961, published a paper entitled "Continent and Ocean Basin Evolution by Spreading of the Sea Floor" introducing the term "sea-floor spreading". Based on the ocean mapping data, Harry H. Hess, in 1962, hypothesized in "History of Ocean Basins," that the earth's outer surface was in motion, causing continents to fragment, move and create new ocean basins. Hess also proposed that new oceanic

STRUCTURE OF THE EARTH Geographic North Pole I I

Geomagnetic North Pole

Geographic Equator

Geomagnetic South Pole

13 Figure 2.3. Earth's magnetic poles. Current magnetic orientation is "normal" with the north and south magnetic poles close to the north and south geographic poles respectively. During periods of reversed polarity, the north and south poles are interchanged.

Geographic South Pole

crust was being formed at mid-ocean ridges by volcanic activity, and destroyed in trenches (Gross 1990). J. Tuzo Wilson in his 1965 paper, "A New Class of Faults and Their Bearing on Continental Drift," elaborated on the ideas, strengthening the concept with the addition of a new type of plate boundary, transform faults. Magnetic Anomalies. Magnetic anomalies on the ocean floor were first noted during submarine detection activities. However, it was not until 1963 that the striped pattern parallel to mid-ocean ridges could be. explained. Drummond Matthews and Fred Vine postulated that the patterns represented reversals in the earth's magnetic field. As new oceanic crust is extruded from mid-ocean ridges, the minerals in the cooling rock are aligned in accordance with the existing magnetic orientation. Each stripe represents a section of the ocean floor formed during a particular magnetic orientation, with the adjacent stripe indicative of a different magnetic orientation or magnetic reversal. At present, the magnetic orientation is "normal" (i.e., the north and south magnetic poles are relatively close to the north and south geographic poles) (fig. 2.3). The.alternating bands of rock with different magnetic orientations create the distinctive "striped" pattern we call magnetic anomalies (fig. 2.4). Sea floor produced now and during other periods of normal magnetic orientation show strong or positive magnetic values. During periods of reversed magnetic fields, the north and south magnetic poles are reversed, and the rocks record a weak or negative pattern. Reversals of the magnetic field occur approximately

14

GEOLOGY

Figure 2.4. Magnetic anomalies. A schematic portrayal of how the distinctive striped pattern of ocean floor magnetic anomalies originates. The sea floor presently being formed along the mid-ocean ridge is of "normal" magnetic orientation. Earlier episodes of normal magnetic fields are marked by the other shaded bands. The white bands indicate sea floor formed during periods of reverse polarity. Note the symmetry of the magnetic anomalies on either side of the mid-ocean ridge. Modified after Ross 1977, and Gross 1990.

every hundred thousand to a few million years. It has been estimated that in the past 76 million years, 171 magnetic field reversals have occurred (Beiser and Krauskoff 1975). Why these reversals take place is not yet known. Not all oceanic crust displays magnetic anomalies. During a period of the earth's history, 80-120 million years ago, magnetic reversals did not occur. As a result, no magnetic anomalies are present in rocks formed during that interval. Deep burial and intense heat may also erase the pattern of magnetic anomalies (Gross 1990). Magnetic anomalies are also a means to map the age and rate of spreading of the ocean floor. By correlating the pattern of oceanic magnetic anomalies with the pattern observed in rocks of known ages on land, we can determine the ages of the various sections of ocean crust, and estimate the rate of sea-floor spreading.

STRUCTURE OF THE EARTH

15

Spreading rates range from less than one centimetre per year on the Mid-Atlantic Ridge, to 16 centimetres per year on the East Pacific Rise. In all cases, the youngest crust is found in a band straddling the spreading ridge, with increasingly older crust on each side of the ridge as the distance from the ridge crest increases. The faster the spreading rate, the thicker the central band of young crust. The oldest oceanic crust, estimated to be 190 million years old, is found in the North Pacific near Asia, and along the margins of the North and South Atlantic. It is believed that approximately half the ocean floor is less than 80 million years old (Gross 1990). Further support for Wegner's Pangaea theory was provided by Sir Edward Bullard in 1965. Using a common depth contour as the edges of all the continents, Bullard was able to piece together the continents to form a supercontinent resembling Pangaea. There were some areas of overlap where relatively new features such as coral reefs and river deltas had developed, but overall the fit was good enough to support Wegner and the Pangaea concept (Gross 1990). Plate Tectonics - The Parts and the Processes. Plate tectonics is the movement of the earth's lithospheric plates (composed of the upper mantle and crust) on the asthenosphere. There are seven major plates and many small plates. The major plates are: Pacific, Indo-Australian, North American, South American, Eurasian, African, and Antarctic. Smaller plates include the Philippine Plate, China Plate, Gorda Plate, Cocos Plate, Nazca Plate, Caribbean Plate, Scotia Plate, Arabian Plate, Iranian Plate, Hellenic Plate, and Juan de Fuca Plate (fig. 2.5). Plate Boundaries. All plates are in contact with several other plates. There are three types of plate boundaries: a) divergent or constructive margins b) convergent or destructive margins c) conservative boundaries or transform faults. Divergent or Constructive Margins. Divergent margins are also known as Atiantic, passive, aseismic or constructive margins. Divergent margins develop when continents rift apart and form new ocean basins. As a result, continental crust and the adjacent oceanic crust are part of the same plate. As the rift widens, the continental margin grows further from the spreading centre and closer to the stable interior. Microcontinents may form if these pieces of continental crust are isolated as a result of rifting or other plate movements (Brown et al. 1989). The actual cause of a divergent margin is not fully understood, but it is thought to begin with the development of crustal stretching, extensional faults, rift basins, and possibly regional uplift and volcanism initiating sea-floor spreading (Hutchinson 1992). A divergent margin, as shown in figure 2.6, generally features the following physical characteristics: a) a continental shelf, gently sloping (average gradient of 0.1°) from the shore and extending to a depth not exceeding 130 m; may be 1500 m wide b) a steep (average gradient of 4°) continental slope, 20 to 100 km wide, continues from the continental shelf to a depth of 4000 to 5000 m, and may feature submarine canyons c) the continental rise, (average gradient of 1 °), may be up to 600 km wide and leads into the abyssal plains (Hutchinson 1992; Brown et al. 1989). Spreading ridges or spreading centres are the site of new crust formation and

Figure 2.5. Tectonic plates of the world. Simplified world map showing the major plates and some minor plates. After Brown et al. 1989, Gross 1990, and Ganeri 1994.

16 GEOLOGY

STRUCTURE OF THE EARTH

17

Figure 2.6. Divergent margins. A generalized cross-section showing the main features of divergent margins: (1) continental shelf, (2) continental slope, (3) continental rise, (4) continental margin, (5) abyssal plain, (6) oceanic ridge. Not to scale. Modified after Brown et al. 1989, and Hutchinson

1992.

occur at divergent or constructive margins. As new crust cools and moves away from the spreading ridge, it increases in density and thickness. New crust is only a few kilometres thick, while old crust reaches thicknesses of 150 km. With increasing density, the crust sinks deeper into the asthenosphere from a depth of 2500 m to 6000 m as it ages (Gross 1990). Two extensively studied examples of early stage continental rifting are the northern Red Sea and the Sal ton Trough in Baja, California (Hutchinson 1992). Active spreading centres in this region are found in the Andaman Sea, Aru Trough, Bismarck Sea, and the Woodlark Basin (Petroconsultants Australasia 1991). Convergent or Destructive Margins. Convergent margins are also known as active, seismic, Pacific-type, or destructive margins. On land, visible signs of convergent margins are active volcanoes, frequent earthquakes, island arcs, and young mountains (Gross 1990). Below the sea surface, subduction trenches mark locations of convergent or destructive margins. Subduction Zones. Most subduction zones involve the subduction of oceanic crust beneath lower density continental crust. Oceanic crust is destroyed as it is drawn into the mantle. Earthquakes occur as the plates drag past each other. Deep- and intermediate-focus earthquakes (100-700 km below the surface) are indicative of subduction zones (Gross 1990). General features of a subduction zone are: trench, accretionary wedge and fore-arc ridge, fore-arc basin, volcanic arc (island arc), and back-arc basin (fig. 2.7). The trench is where the descending plate first contacts the material from the overriding plate as it heads towards the mantle. The accretionary wedge forms when sediments and other materials are scraped from the subducting plate and accumulate (like the action of a snowplow pushing the snow ahead to clear a path), at the front of the overriding plate (Hamilton 1989). It is composed of deformed rocks. The fore-arc ridge is the summit of the accretionary wedge. The fore-arc basin is found landward of the accretionary wedge and contains

18

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Figure 2.7. General features of a convergent margin are: (1) active trench, (2) accretionary wedge, (3) fore-arc ridge, (4) fore-arc basin, (5) frontal arc, (6) volcanic arc, (7) back-arc region. Modified after Curray 1989, and Hutchinson

1992.

relatively less-deformed sediments (Hutchinson 1992). The frontal arc located between the accretionary wedge and the volcanic arc is a zone of uplift and deformation. The volcanic arc is a zone of active igneous activity (Hutchinson 1992). The back-arc region is found behind the volcanic arc and may contain marginal (back-arc) basins or ancient, inactive arcs. The Benioff zone, also known as the Benioff seismic zone, is a seismic zone where earthquake foci cluster. An earthquake focus is the point within the earth which is the centre of an earthquake (Bates and Jackson 1980). The Benioff zone stretches downward from the oceanic trench, dipping toward the continents usually at an angle of 45°. The Scenario Involving Old Subducting Crust. When the subducting or descending oceanic plate is old and therefore dense, it descends as a steeply dipping slab. This is a common scenario in the western Pacific basin where the oldest oceanic crust occurs (fig. 2.8). Subduction of the oceanic crust may also be accompanied by back-arc spreading and basin formation as is the case in the area of the Mariana Trench (Gross 1990). According to Hutchinson (1992), subduction involving old oceanic crust is "low-stress", and commonly features a small accretionary wedge, few large earthquakes, igneous rocks with a narrow basaltic compositional range, a steeply dipping descending slab, and a well-developed back-arc basin. The Scenario Involving Young Subducting Crust. When the subducting plate is composed of relatively young, still-buoyant crust, it descends as a shallow-dipping slab (fig. 2.9). The east coast of the Pacific (i.e., the west coast of North and South

STRUCTURE OF THE EARTH

19

Figure 2.8. Subduction involving old oceanic crust. The main features to note are: (1) steeply dipping oceanic plate, (2) deep trough with little accumulation of sediments or a small accretionary wedge on the upper plate, (3) active island arc, (4) well-developed back-arc basin and, (5) extinct arc. (Few large earthquakes.) Modified after Gross 1990, Curray 1989, and Hutchinson

1992.

America) is an example of where this type of shallow subduction is taking place. Active volcanoes and young mountains line the coast. The mountains are formed from material scraped off the subducting plate (Gross 1990). Subduction involving young oceanic crust is classified by Hutchinson (1992) as "high-stress", and features a large accretionary prism, large shallow earthquakes, igneous rocks of varied composition, and a shallow-dipping plate. When material is scraped off an oceanic plate and accretes to a continental plate, it is known as an ophiolite. Ophiolites are a specific type of exotic terrane which is any fragment of continental or oceanic plate welded onto a continent. Ophiolites are often rich sources of sulphide minerals formed at spreading centres, and are mined for copper, lead, zinc, and silver (Gross 1990). Subduction is presentiy occurring off the coasts of Sumatra, Java, along the East Sunda Arc, Banda, Seram, Sangihe, North Sulawesi, Cotabato/West Sangihe, and Halmahera (Petroconsultants Australasia 1991). The trenches where subduction is taking place are known as active trenches, as opposed to inactive trenches where subduction has ceased. Collision of Two Continental Plates. Collision of two continental plates, does not produce a subduction scenario because both plates are composed of fairly buoyant crust. Instead, overriding and uplift can occur with one plate being folded and thrust upon the other (Gross 1990). This is what happened when the Indian Plate collided with the Eurasian Plate resulting in the formation of the Himalaya Mountains. Collisions are also taking place in Sulawesi, and Timor-Tanimbar (Petroconsultants Australasia 1991). Subduction Involving Two Oceanic Plates. When one oceanic plate is subducted

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Figure 2.9. Subduction involving young oceanic crust. The main features to note are: (1) shallowly-dipping oceanic plate, (2) relatively shallow trench, (3) large accretionary wedge, (4) forearc basin, (5) active volcanoes. Earthquakes are generally large and shallow. Modified after Gross 1990, and Hutchinson

1992.

beneath another oceanic plate, a trench is present, and a volcanic island arc forms on the overriding plate. On the opposite side of the island arc, away from the trench, commonly a marginal or back-arc basin forms as a result of sea-floor spreading. Conservative Boundaries or Transform Faults. The third type of plate boundary involves lateral movement with transform faults (fig. 2.10). Transform faults are a type of strike-slip fault in which the plates slide horizontally past each other with oceanic crust neither created nor destroyed (Roberts 1989). Many transform faults are associated with mid-ocean ridges and are seismically active. They run perpendicular to the line of the ridge, offsetting the ridge, and terminate where they meet the ends of the two offset ridge segments which they connect (Brown et al. 1989). The extension of a transform fault beyond those points is known as a. fracture zone, and is characterized by no relative sideways motion, and low seismicity (i.e., small earthquakes) as it is located within a single plate. Examples of transform faults in Indonesia are in Sumatra and Sorong (Petroconsultants Australasia 1991). One of the world's best-known examples of a transform fault is the San Andreas Fault in California, U.S A.

STRUCTURE OF THE EARTH

21

Figure 2.10. Transform faults. The ridges of the earth are marked by a series of transform faults which offset the ridges resulting in the irregular outlines. Movement along the faults during seafloor spreading produces shallow earthquakes marking the active section of the fracture zone. Beyond the active area the fracture zone features steep ridges and valleys. (1) Mid-Atlantic Ridge, (2) Southwest Indian Ridge, (3) Indian Ridge, (4) Southeast Indian Ridge, (5) Ninety-east Ridge, (6) Java Trench. Modified after Ganeri 1994, Earth 1992, and Gross 1990.

Earthquakes Rocks, like people, can be affected by stress. Stress builds up in rocks as lithospheric plates move past each other. When the stress is greater than the strength of the rocks, the rocks are strained, then fail (i.e., fracture along zones of weakness), faults are formed, and the energy is released in the form of an earthquake (Keller 1979). Although plates are often described as "sliding" past another plate, the movement is not smooth and there is much friction where the plates meet. Movement along faults produces seismic waves to depths of 700 km (Ross 1977), and in the near-surface of the earth. An earthquake may be described as a sudden motion of the earth caused by faulting or volcanic activity. As many as a million earthquakes each year are recorded on seismographs, with most escaping the notice of human populations. The centre of an earthquake is called the focus or hypocentre. The focus is the point deep within the earth of the initial rupture and where the strain energy is first converted to elastic wave energy (Bates and Jackson 1980). Shallow-focus earthquakes are those with a focus at 70 km or shallower. Deep-focus earthquakes begin below 300 km (Kingston 1988). The point on the earth's surface above the focus is called the epicentre. The epicentre is generally located directly over the focus, except when the earthquake originates deep within a subduction zone.

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The Mercalli Intensity Scale. The Mercalli intensity scale, devised in 1902, is an arbitrary scale of earthquake intensity, ranging from I (detected only by instruments) to XII (almost total destruction). Its adaptation to urban conditions is known as the modified Mercalli scale (Bates and Jackson 1980). Modified Mercalli scale (from Whitten 1972, and Spencer 1972): I. Instrumental. Detected only by seismographs. II. Feeble. Noticed by only a few persons at rest. Delicately suspended objects may swing. III. Slight. Resembles vibrations caused by heavy vehicle traffic. IV. Moderate. Felt by most people. May waken some. Rocking of free-standing objects. Walls crack. Sensation like heavy truck striking building. V. Rather strong. Sleepers awakened. Bells ring. Widely felt. VI. Strong. Felt by all. Trees sway. Some damage from overturning and falling of objects. VII. Very strong. General alarm. Everyone runs outdoors. Damage is negligible in buildings of good design and construction. Slight to moderate damage in well-built ordinary structures. VIII. Destructive. Fall of factory smoke stacks, columns, monuments, and walls. Heavy furniture overturned. IX. Ruinous. Ground begins to crack. Houses collapse. Underground pipes break. X. Disastrous. Ground badly cracked. Many buildings destroyed. Some landslides. Rails bent. Water splashed over banks. XI. Very disastrous. Few buildings remain standing. Bridges and railways destroyed. Broad fissures in the ground. XII. Catastrophic. Total destruction. Waves seen on ground surface. Objects thrown into the air. The modified Mercalli intensity scale is based on effects which can be observed, and refers to the "violence of the earthquake motion" (Spencer 1972). However, this scale is difficult to apply for accurate, quantitative, worldwide comparisons of earthquakes. In 1935, the Richter magnitude scale was devised. The Richter Magnitude Scale. The Richter magnitude scale, more commonly used today, and generally referred to as the Richter scale, is a logarithmic scale, based on the amount of energy released at the focus, as recorded by a seismograph. A seismograph records earthquake waves. The amplitude of the largest wave determines the magnitude (Keller 1979). An earthquake of magnitude 6 produces a displacement on a seismograph 10 times larger than does a magnitude of 5. A magnitude 5 earthquake releases approximately 1021 ergs of energy which is equivalent to the first atomic bomb detonated in 1945, or 20,000 tons of TNT (Spencer 1972). However, although the total energy released may be comparable, its mode of dispersal produces very different effects depending on whether it was released in a highly concentrated form as in the atomic bomb, or widely dispersed as in an earthquake. Table 2.1 provides a qualitative description of the Richter scale. Effects of Earthquakes. In addition to the well-known disastrous effects of earthquakes on population centres, earthquakes affect marine environments (see section on coral reefs and natural disturbances). The primary effects of earthquakes are vio-

STRUCTURE OF THE EARTH

23

lent ground/substrate motion, surface rupture, and permanent substrate displacement of a metre or more (Keller 1979). Secondary effects maybe divided into short-term events such as landslides, tsunamis and floods. Long-range effects are regional subsidence or emergence of landmasses. Distribution of Earthquakes. Active seismicity marks plate boundaries. The zones of seismicity tend to be narrow when associated with mid-oceanic spreading centres, and strike-slip faults. Wider zones occur above subducting plates, and in those parts of continents undergoing distributed extensional, strike-slip, and compressional deformation (Hamilton 1979). Earthquakes of a magnitude greater than 8.0 occur primarily along subducting plate boundaries, and less frequently in continental strike-slip and compressional deformation situations. Distribution of Earthquakes Worldwide. The earth can be divided into 10 regions based on seismic activity: 1. The circum-Pacific belt contains most of the shallow- and intermediatedepth earthquakes, and almost all deep-focus earthquakes; 2. The Alpine belt of Europe and Asia, containing the Alps and Himalayas, is the other significant zone of shallow and intermediate-depth earthquakes; 3. The Pamir-Baikal zone of central Asia; 4. The Atlantic-Arctic belt; 5. The belt of the central Indian Ocean; 6. Rift zones, notably those of east Africa; 7. A wide triangular active area in eastern Asia, between the Alpine belt and the Pamir-Baikal zone; 8. Minor seismic areas, usually in regions of older mountain building; 9. The central basin of the northern Pacific Ocean; almost nonseismic except for the Hawaiian islands; 10. The stable central shields of the continents, also nearly nonseismic. Distribution of Earthquakes in Indonesia. Indonesia, located within the famed "Ring of Fire," is frequently hit by earthquakes covering a range of magnitudes. Approximately 10% of the world's seismicity occurs in the Indonesian Archipelago. Katili (1985) gives a brief summary of the distribution of Indonesia's shallow, interme-

Table 2 . 1 . Scales of magnitude on the Richter scale. Magnitude

Description

2.5 > 4.5 6.0 - 7.0 7.0 - 7.7 7.7 - 8.6

Just large enough to be felt Capable of causing very local damage Potentially destructive Major earthquake Great earthquake

Source: Leet and Judson 1965.

focus > 100 km

After Katili 1985, and Hamilton 1974.

Figure 2.11. Distribution of shallow and deep earthquake epicentres in Indonesia.

ocus < 100 km

:

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25

diate, and deep earthquakes (fig. 2.11). Indonesia's shallow-focus earthquakes most commonly occur above the subducting plate boundaries of the Java-Sumatra Trench and the Banda Trench, within the active collision zone between Sulawesi and Halmahera, and associated with transcurrent faults (i.e., the Great Sumatran fault zone, the Palu-Koro-Matano fault zone, the Gorontalo fault zone, and the Sorong fault zone). The intermediate-depth (focus depth of 100-300 km) earthquakes are generally spread along "the whole volcanic arc divided in the middle by the axis of active volcanoes in Sumatra, Java, the Lesser Sunda Islands, Sulawesi and Halmahera" (Katili 1985). The deep earthquakes (focus depths of 500-800 km) are clustered along an east-west trending belt from the Java Sea to the Banda Sea, and a north-south trending belt from Sulawesi to Mindanao.

VOLCANOES To many, the mere mention of "Indonesia" brings to mind the archipelagic nation's most famous volcanoes, Krakatau and Tambora. Krakatau, immortalized by Hollywood's movie industry, although geographically misplaced in the film Krakatau, East ofJava, ranked fourth in the world's greatest historic eruptions (Hutchison 1982). The 1815 eruption of Tambora on the island of Sumbawa was the most violent explosion of recorded history, yet is relatively unknown outside geological circles. Indonesia's two most famous eruptions will be discussed in greater detail after an introduction to volcanoes and volcanic activity. The Shapes of Volcanoes Volcanoes come in several shapes and sizes. "Volcanoes show a wide variety of forms, depending largely upon the composition of the erupted material and hence the style of eruption" (Thorpe and Brown 1985). The three main shapes are cinder cones, shield volcanoes and composite or stratovolcanoes. Several factors determine the shape of a volcano: • Land surface; • Type and nature of material ejected (i.e., viscous lava and cinder build up steep cones, while more fluid lava flows further away from the vent and results in wide-based mountains); • Forcefulness of the eruption; • Duration of the activity; • Eruptive history of the volcano. Cinder Cones. Cinder cones are built when pyroclastics, such as cinders, and ash, pile up around the vent. A crater normally surrounds the vent. The slopes are generally steep (greater than 10°) and symmetrical. These volcanoes are usually basaltic or andesitic. The magma contains a relatively high gas content, resulting in higher explosivity (van Bemmelen 1949) than cumulo-volcanoes (volcanic domes) or lava shields (shield volcanoes of the basaltic type). Examples: Lamongan, East Java and Vesuvius, Italy.

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Shield Volcanoes. Shield volcanoes feature the largest cones. This type of volcano has a broad, shield, or low-slope profile. Its diameter may be between 100 and 200 km (Thorpe and Brown 1985). Shield volcanoes represent the largest discrete volcanic form, but they are rarely preserved in the geological record, since they occur mainly as oceanic islands (Thorpe and Brown 1985). They are produced by eruption of low-viscosity lava from either a central vent, fissures or parasitic vents. The flow generally consists of very fluid basaltic lava or rhyolitic ash flows (Bates and Jackson 1980). The gas content of the magma is low. They are also sometimes referred to as lava domes. Examples: Sukadana, South Sumatra and Mauna Loa, Hawaii. Composite Cones (Stratovolcanoes). Stratovolcanoes, or composite cones, have a central vent from which lavas as well as pyroclastics are expelled. This type of volcano is composed of alternating layers of lava and pyroclastics, resulting from the prolonged activity of a central vent causing the formation of bedded-volcanoes (van Bemmelenl949). Stratovolcanoes are the most common shape for andesitic volcanoes. They have steep, irregular conical forms with diameters of 10-40 km. The volcano often has a shape combining cinder and shield volcanoes. There are often many dikes and sills. The lava, viscous and acidic, is generally restricted to the area of the volcano, but the pyroclastic material may be transported by wind more than 1000 km (Thorpe and Brown 1985). These deposits are often used by stratigraphers as markers or reference points. Examples: Merapi, Central Java and Fujiyama, Japan. Composite volcanoes clearly show a zonation of volcanic products which may be divided into the central, proximal arid distal zones with increasing distance from the central vent (Thorpe and Brown 1985). The central zone is located within 2 km of the central vent. It is characterized by lava conduits. The volcanic products associated with this zone are coarse, poorlysorted pyroclastic materials which have been deposited close to the vent. The proximal zone, situated 5-15 km from the central vent, contains a higher proportion of lava flows and a variety of pyroclastic flow deposits. The distal zone is found beyond the proximal zone, and consists of pyroclastic flow deposits associated with fine air-fall deposits dispersed by the wind away from the volcano. Other Volcanic Landf orms Caldera. A caldera is a large, basin-shaped volcanic depression, more or less circular, the diameter of which is many times greater than that of the included vent or vents (Bates and Jackson 1980). There are two main types of volcanic calderas depending on the mode of formation (i.e., collapse caldera, and explosion caldera). A collapse caldera is formed when the top of the magma chamber collapses into a void or cavity created by the removal of magma either by large volume eruptions of lava or pyroclastics, by subterranean withdrawal of magma, or contraction of magma as it cools and crystallizes. The top of the volcano collapses into the void (subsidence phenomena) forming an enclosed or partially enclosed depression called a caldera which maybe tens of kilometres across. Most calderas are of this type (Bates and Jackson 1980). An explosion caldera is formed by the explo-

VOLCANOES

27

sive removal of the upper part of a volcanic cone. This type of caldera is extremely rare, and is small in size according to Bates and Jackson (1980). Plateau Basalts. Plateau basalts are the most extensive volcanic landform, with the basaltic lavas covering areas up to 105 km2 (Thorpe and Brown 1985). The lava originates from cracks or fissures (i.e., fissure eruptions) rather than from a central vent (Rhodes 1991). Plateau basalts erupted in rapid succession over vast areas and have, at times, flooded sectors of the earth's surface on a regional scale (Bates and Jackson 1980). Examples are known from India, Iceland and the United States of America where the Columbia River Plateau has an estimated volume of 417,000 km covering an area greater than 250,000 km (Spencer 1972). These lavas range in age from Eocene (40-60 million years old) to Pleistocene and Recent (i.e., less than 2 million years old). Volcanic Domes. Volcanic domes are dome-shaped, bulbous masses of hardened lava. They form above and around volcanic vents. When a volcanic dome forms on the side of, or close to a larger volcanic cone, it is known as a parasitic volcano. Parasitic volcanoes are generally 1-2 km diameter, created by a single, short-lived eruption (Thorpe and Brown 1985). Lahar. Lahar is an Indonesian word which has been adopted by geologists worldwide to describe mudflows which contain debris and angular blocks mostly of volcanic origin (van Bemmelen 1949). The transporting medium is a mixture of cool (60 km) that they are situated in mostly oceanic conditions. However, the numerous patch reefs and extensive shallow-water platform reefs (e.g., P. Panjang with a 143 km reef flat) that comprise the barrier reef system are under the direct influence of the Berau River as well as the Bulungan River, 60 km to the north. The sediments of the inshore fringing reefs contain relatively high percentages of fluvial deposits. The reef flats exhibit distinct lagoonal characteristics, supporting high abundances of burrowing macroinvertebrates (fig. 4.9). The reef flats are extremely productive coral reef habitats, and even with high percentages of fluvial deposits, they clearly are of reefal origin and should be classified as such. On many volcanic islands (e.g., Solor, Adonara, Lembata) the reefal carbonates may constitute less that 50% of the coastal sediments, even though the entire island may be fringed by a reef. In other areas, sediments along large sections of coastline may be mostly of non-reefal origin (e.g., Merdeka Bay, Lembata Island; Luwuk, Central Sulawesi), even though well-developed fringing reefs are present. Taking these, two criteria (i.e., light and sediments) under consideration (the third criteria stipulating >50% dominance by tropical organisms is not relevant here), Crossland et al. (1991) delineated the outer boundary of coral reefs as:

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133

Figure 4.9. Subtidal reef flats of inshore coral reefs in the Berau Islands are highly productive habitats with distinct lagoonal characteristics, resulting from the relatively sheltered position of the reef complex, and its close proximity to a major river delta. Raburabu patch reef at high spring tide, East Kalimantan. Photo by Tomas and Anmarie

Tomascik.

".. .the bottom of the euphotic zone or transition to 100 km2, and no basement). Subsequent subduction of the region resulted in transgression and major carbonate deposition. Following the formation of the carbonates, both reefal and non-reefal, in late Miocene, sea levels dropped considerably thus exposing the limestones. As a result, the surface topography of the southeast Java mountains has been heavily karstified under the onslaught of a humid tropical climate. Similar heavily karstified limestone deposits occur along the south coast of Central Java (i.e., Gunung Sewu range - fig. 5.25). The largest, and the most studied Miocene carbonate deposits are the Wonosari limestones, found in the Wonosari district. The Gunung Sawu reef-limestones have a thickness of about 800 m (van Bemmelen 1949). Uplift, tilting and block faulting of these carbonate deposits in the middle Pleistocene resulted in extensive karstification. Stratigraphic studies indicate that Wonosari limestones were at sea level during most of the Neogene. Along the north coast of Central and East Java, major limestone deposits were laid down in the Rembang beds, a narrow mountain chain running roughly from Purwodadi to Tuban. These oil- and coal-bearing deposits are economically significant. Early Miocene coral fossils have been found in relatively rich reefal deposits. According to Umbgrove (1946c), about 17% of recent coral and molluscan fauna is represented in the early Miocene (about 20 Ma B.P.) Rembang deposits, while a number of coralliferous deposits from the Oligocene of the Indonesian Archipelago contained no extant species (see table 5.1). During Leg 5 of the Snellius-II Expedition to Linta Strait and northeastern Komodo, two Miocene reef deposits were discovered on two volcanic islands, Pulau Sabita and Gili Lawa Laut (fig. 5.26). Table 5.1 indicates that the Miocene scleractinian fauna of this region was similar to present-day communities; however, the absence of Acropora has been noted as remarkable (van der Land and Sukarno 1986). The two fossil fauna, with a total of 23 genera, are very similar in character, suggesting that they may have been from the same period, but not older than early Miocene as is indicated by the presence of Seriatopora. The Sabita deposit also contained interesting fossils which apparendy resembled Mesozoic stromatoporoids (sclerosponges) (van der Land and Sukarno 1986).

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Young volcanoes Subrecent basal flows (N. of Solo) Older volcanoes, high terraces J Kabuh Beds Puljangan Beds (black clays, volcanic breccias, etc.) Damar Beds (Ungarea area)

Banjah Beds, Oje Beds (Southern Mts.) Basal layers of the Sentale Beds (West-Proge Mts.) I Kerak Beds (Kendeng Zone) | Merawu Beds (West of Ungaran) 1 Djanggrargan Beds (West-Proge Mts.) Old Andesite formation (Southern Mts.) Penjaten series (West of Ungaren) Old Andesite formation (West-Proge Mts.)

I

j Lutul Beds (West of Ungaren) i Pelang Beds (near Djuwangi)

I Eocene (Djiwa, Gomping, Godean) ] Pretertiary (Djiwe Hills)

Figure 5.25. Map of the south coast of Central Java showing location of limestone formations, highlighted by a heavy line. From van Bemmelen

1949.

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195

Figure 5.26. Map of east Komodo showing the locations (stars) of Pulau Sabita in Linta Strait and Gili Lawa Laut off the northeast coast of Komodo. From van der Land and Sukarno 1986, p.3-6; Leg 5.

Further to the northwest, reefs and carbonate platforms were deposited along the margins of the South China Sea. In comparison, deep-water clastic sediments were deposited in the Gorontalo and Bone Basins. Further to the west, a massive carbonate platform extended from southern Sulawesi to the western side of the south Makassar Basin. More importantly, however, is that reefal build-ups and associated carbonate deposition formed thick reefal units, many of which may be important oil reservoirs (Haq et al. 1987). The late Oligocene to middle Miocene transgression was associated with climatic warming (Savin et al. 1985), which resulted in widespread latitudinal expansion of coral reefs (i.e., from Japan to New Zealand), mainly as a result of a reduced oceanic latitudinal temperature gradient (Fulthrope and Schlanger 1989). From the middle Miocene high sea-level stand, global sea levels dropped in three progressive stages, reaching a low of about 200 m below present-day sometime during the end of the late Miocene. From the late Miocene to Quaternary, global sea levels experienced a number of cycles of rapidly fluctuating sea levels. The most impressive rise in sea level occurred at the Miocene/Pliocene boundary (about 5 Ma B.P.) when sea levels reached 140 m above present (Hutchison 1989). Early Pliocene was a period of large-scale carbonate deposition along the northwest

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shelf of Australia, as well as a period of active tectonism throughout eastern Indonesia. The central uplands of Papua New Guinea as well as the Tamrau Mountains of Doberai Peninsula, Irian Jaya, were uplifted by tectonism. Fossil-rich late Neogene carbonate deposits abound in the Van Rees Mountain range of northern Irian Jaya. Some of these have shed some light on the longevity and persistence of some coral species groups. For example, Umbgrove (1946c) found that about 66% of fossil corals in the early Pliocene deposits of the Van Rees Mountains belong to extant species (see table 5.1). Just to the west, Seram and Buru Islands were partially uplifted as a result of a dextral shear along the Seram Trough. Thick, deep-sea sediments were deposited throughout the Moluccas Sea. The Sula Platform (i.e., the Banggai-Sula microcontinent), which is considered to be a fragment of the North Australia-New Guinea passive continental margin, collided with Sulawesi sometime in the late Neogene (Klompe 1956; Hamilton 1979; McCaffrey et al. 1981; Pigram et al. 1985; Garrard et al. 1988; Davies 1990). As a result, the eastern part of Central Sulawesi (e.g., Luwuk) was subsequendy uplifted, creating a series of uplifted coral reef terraces that reach an altitude of about 400 m above present-day sea level (Sumosusastro et al. 1989). Using U / T h dating techniques, and a sea level curve from Huon Peninsula, Papua New Guinea, Sumosusastro et al. (1989) were able to determine that the highest coral terrace (410 m)

Table 5.1. Miocene scleractinian genera from Pulau Sabita and Gili Lawa Laut, Lesser Sunda Islands. Numbers indicate number of samples collected. See figure 5.26 for station location. Scleractinian genera Acanthastrea Goniastrea Pontes Astreopora Galaxea Montastrea Symphyllia Fawa Goniopora Seriatopora Cyphastrea Leptoria Pavona Favites Platygyra Echinopora Lobophyllia Pocillopora Leptastrea Stylocoeniella Diploastrea Oulophyllia Pleisiastrea

Pulau Sabita 10 4 7 9 2 1 8 6 5 1 1 1 1 22 12 3 1 2

Data from van der Land and Sukarno 1986.

Gili Lawa Laut 22 4 11 2

7 6 1 4 3 2 2 1 1 1 1

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197

was uplifted at a rate of about 184 cm.ka 1 . Deep-water elastics derived from the uplifted deposits in North Sulawesi were deposited in the Gorontalo Basin. To the south, the Buton continental fragment collided with the southeastern arm of Sulawesi, which subsequently raised Pleistocene coral reefs up to 703 m above sea level at the southern part of the island (i.e., Mt. Kontu) (van Bemmelen 1949; Hutchison 1989). During the early Pliocene, Java was an area of marine deposition; however, uplifting did not occur. To the north, the present location of the South China Sea became established, but most coral reefs in the East Natuna Basin were drowned by the rapidly rising sea levels. From a global perspective, the most important event that took place during the mid-Pliocene was the closure of the Central American Seaway and the rise of the Panamanian land bridge (about 3 Ma B.P.). This event effectively isolated the Caribbean Province from the Indo-Pacific, which greatly influenced present-day biogeographical patterns. The Quaternary Wide and rapid eustatic sea-level fluctuations that followed in the Pleistocene were all associated with the growth and melting of polar and continental glaciers. The Pleistocene is generally known as the Ice Age period. Associated with the wild swings in climatic conditions were corresponding changes in ocean circulation patterns. However, the development of modern reefs has been mainly influenced by recent sea-level fluctuations, especially since about 120,000 years ago. Coral stratigraphy of the Tertiary and Quaternary of Indonesia has been extensively studied since the 1920s (Umbgrove 1924, 1926, 1938, 1939b, 1942, 1943a,b, 1945, 1946a,b; Osberger 1956). Based on the extensive fossil record, Umbgrove (1946c) was able to show that since the Miocene there was a progressive increase in the percentage of Recent (Holocene) scleractinian species (table 5.2). The pattern in table 5.2 has shown to be applicable in a global context, and we now know that many of the extant scleractinian coral genera have survived since at least the early Tertiary. The table illustrates that scleractinian species in Indonesia have undergone major changes since the late Miocene; however, the work by Veron (1995) shows that at the generic global level, changes were relatively minor. The data presented in table 5.2 suggest that in spite of great climatic upheavals, and wide eustatic sea-level fluctuations during most of the Plio-Pleistocene, scleractinian corals survived without major extinctions in their generic ranks (Wells 1956; Pauly 1991; Buddemeier 1993). However, at the species level major changes have occurred within the past few million years. None of the extant Indonesian species have a fossil record beyond the late Oligocene. The latest drop in sea level (i.e., 135 m below present) that occurred about 18,000 years ago apparently had very little impact on scleractinian diversity, but it had a major impact on present-day biogeographical distribution patterns. The Indonesian and Philippine Archipelagoes became refuge places, from which adjacent regions, such as the Great Barrier Reef, became repopulated. The post-glaciation Indonesian/Philippine source pool may explain, at least partly, the surprising absence of scleractinian endemism in the region (Veron 1995). However, C. Wallace has recently found a number of new endemic Acropora species in various parts of the archipelago (C. Wallace, pers. comm.).

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Potts (1983) suggested that the lack of endemism in the Indo-Pacific region is related to reduced rates of speciation during periods of rapidly fluctuating sea levels that were common during the Plio-Pleistocene. He proposed that during these environmentally stressful times, corals did not have sufficient time for speciation, but instead they maximized their intraspecific variability, which resulted in extensive variation in external morphology (e.g., Acropora and Montipora). He also suggested that since the late Pliocene, speciation in corals has been suppressed, at least in the Indo-Pacific region. This, however, seems contrary to recent studies by Boekschoten et al. (1989) who, based on their Indonesian field data, concluded that, at least for Acropora, evolutionary stasis has not set in. An alternate theory by McManus (1985), however, suggests that the rapid sea-level fluctuations did not retard, but rather facilitated, speciation on both land and sea. A full knowledge of Indonesian endemics would go a long way to reconcile these different viewpoints; unfortunately, very little can be said of Indonesian endemics since there is a serious lack of quantitative data on the subject. While some information exists on the distribution of the early scleractinian communities (Gerth 1923; 1925,1932; Umbgrove 1924,1926, 1946a, 1946b; Oosterbaan 1985), there is little information on zonation and community structure of these early coral reef communities. Boekschoten et al. (1989) were among the first to gather paleontological evidence which suggested that the generic composition of Indonesian coral reefs has remained relatively unchanged during the past 25 million years. Indeed, present-day coral reef communities are most likely very similar in terms of community structure and function to coral communities that existed during the interglacial periods or sea-level stillstands during the Plio-Pleistocene (Jackson 1992). In general, the Pliocene scleractinian deposits in Indonesia contain mostly extinct species belonging to extant genera (Boekschoten et al. 1989). Nevertheless, there are some coral species that show remarkable persistence in maintaining their general morphology. For example, Boekschoten et al. (1989) found a Pliocene Pachyseris speciosa (Dana) in a reef talus deposit on Guang Island, southwest Selayar, that is identical to the extant species (i.e., the species has maintained its

Table 5.2. Percentage of Recent (Holocene) scleractinian coral species in the fossil record since the early Eocene (57 Ma B.P.). 'Ma B.P.' refers to million years before present. Epoch Pleistocene - Holocene Pliocene Late Miocene Middle Miocene Early Miocene Oligocene Eocene

Time period* (Ma B.P.) 1 - Recent 5-1 11-5 17-11 24-17 37-24 5 7 - 37

Percent of Recent scleractinian species 70-100 50-70 30-50 10-30 0-10 0- ? 0

* These are estimates only corresponding roughly to the Letter-classification as given by van Bemmelen (1949), and should be interpreted with caution. Based on Umbgrove 1946; van Bemmelen 1949.

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199

Box 5.2. Evolution of Indonesian coral reefs. G.J. Boekschoten; Earth Sciences, Free University Amsterdam; M. Borel Best; National Museum Natural History, Leiden. The Indonesian Archipelago is a result of a collision of three major tectonic plates (IndoAustralian Plate, Pacific Plate and Eurasian Plate) during the Tertiary. Prior to the collision, each of these three crustal elements had their own unique marine biotic identity. The collision resulted in a wide range of passive and active underwater slopes, consisting of older rocks and newer sediments or lava streams. The variety of provenance of the reef organisms, and the great range of ecological opportunities offered by the complicated ranges of island arcs and continental fragments, may explain the unique species richness of the eastern Indonesian coral reefs. The Asian portion of the western part of the Indonesian Archipelago (i.e., Sunda Shelf) had little coral reef development. However, fringing and barrier reefs developed around the seaward edge of the shelf. During the low sea-level stands in the Tertiary and Quaternary, many of the coral reefs fringing the Sunda Shelf were smothered in mud and sand, and some barrier reefs drowned. ( In contrast, coral reefs flourished in the Pacific part of the Indonesian Archipelago, the Luzon-Kalimantan-West Irian triangle. Many coral reefs in eastern Indonesia developed on topographically steep ridges that supported high species diversity. These submarine ridges may have been important refuges for corals and other reef organisms during the Quaternary, characterized by wide sea-level fluctuations. The islands in this region did not have much upward movement, and, as a result, fossil reefs and terraces that are characteristic of southern Indonesia were not formed. In the Australian portion, technically the most active, of the Indonesian Archipelago (southern Indonesia, Moluccas, western part of Sulawesi and lower Sunda islands), a variety of coral reefs (e.g., fringing, patch, barrier reefs and atolls) developed. Tectonic events associated with the rapid northward motion of the Sahul Shelf (Northern Australia) towards the Eurasian Plate resulted in many elevated reefs and reef terraces (fig. 5.27). Fossil reef limestones occupy a considerable surface portion of many Indonesian islands (fig. 5.28). The limestone ranges along the southern coastline of Java and the Lesser Sunda Islands consist partly of this reefal material. Indonesia's considerable oil reserves are held in the subsurface of fossil Tertiary reef rock. These limestone outcrops are also an important source of "marble" slabs embellishing banks and shop fronts. The polished surfaces of most marble contain many lentilshaped shells of fossil foraminifera (Lepidocyclina a.o.), reef-dwelling benthic organisms characteristic of the middle Tertiary reef assemblages (fig. 5.29). Foraminifera became less important in younger Tertiary reef assemblages. Sea-level changes associated with the growth and melting-down of polar ice caps became more important towards the Pleistocene Ice Ages. A low sea-level stand of at least 110 m below present-day sea level still existed some 14,000 years ago. The sea level rose rapidly to about the present level afterwards. The coral reef zone must have migrated up and down several times during the Pleistocene within this range. The remarkable paucity of the genus Acropora in Pleistocene reefs has been linked with this process. The preference of Acropora species for reef flats and upper reef slopes, areas that were first destroyed during the transitional phases of rapid transgression, may have been the key factor responsible for their low abundance during the Pleistocene. The spasmodic history of the Pleistocene sea level resulted in the development of reef terraces along the coasts, that were also shaped by local tectonic uplifts. The sea level during the last 10,000 years did not remain stable, but fluctuated mainly as a result of changes in

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Box 5.2. (Continued.)

Figure 5.27. The elevated reef terraces of Binongko Island, Tukang Besi Islands, Southeast Sulawesi. Note an unusual barrier reef at centre right.

Figure 5.28. Uplifted Pleistocene reef limestones on Sumba Island, East Nusa Tenggara.

CENOZOIC REEFS

201

Box 5.2. (Continued.)

Figure 5.29. Fossil foraminifera Lepidocyclina (Foraminiferida) in Miocene limestone deposits. Figure 5.30. Coral cays of the Spermonde Archipelago, South Sulawesi.

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GEOLOGICAL HISTORY OF REEFS

Box 5.2. (Continued.) wind and current patterns. Present sea-level stand, in much of Indonesia, is about 2-3 m below the highest Holocene sea-level stand. Consequently, many sandy reef flats have emerged as coral cays with low flat surfaces. The Spermonde Archipelago in southwest Sulawesi is a good example (fig. 5.30); the Kepulauan Seribu, northwest of Jakarta, is another. The latter possibly reflects a pre-existent pattern of Pleistocene hills that were covered by coral reefs since the onset of the Holocene.

form for the past 5 million years). Guang Island (fig. 5.31) is a part of a small horst, running parallel with the Selayar horst, whose geological setting and coral fauna suggest Pliocene age (van der Land and Sukarno 1986). A total of 23 scleractinian genera (table 5.3) were found in the outcrop, including specimens of Tubipora, which has previously not been described as a fossil in the scientific literature. Table 5.3 illustrates that the generic diversity of the Pliocene reefs at Guang and Bahuluang Islands was remarkably similar to present-day reefs, even though specieslevel taxonomy was most likely much different. The most frequently found specimens were Goniastrea (16), Acropora (12) and Pontes (10) (table 5.3). The abundance of these large massive genera may be a reflection of their greater resistance to weathering rather than of ecological significance. The Guang outcrop also contained Pecten (Pectinidae) and Tridacna (Tridacnidae) fossils. Another limestone outcrop bearing fossil fauna of the Pliocene was found on Bahuluang Island, which is located just to the south of Guang Island. According to van der Land and Sukarno (1986), the limestone outcrop on Bahuluang Island belongs to the Selayar member of the Walanae formation. While the fossil fauna of the Bahuluang outcrop was similar to Guang, it was much less extensive. Studies of the Plio-Pleistocene and early Pleistocene fossil reefs in Indonesia revealed a conspicuous absence of Acropora from die coral communities. Table 5.4 demonstrates the species composition of Plio-Pleistocene reefs in Nias, off the west coast of Sumatra and Pleistocene communities of Sumba. The absences of Acropora and Montipora are striking, since Acropora is considered the most important reef-builder in recent Indonesian reefs, as well as the most diverse genus (Moll 1983). According to Boekschoten et al. (1989), the total absence of Acropora from Indonesian Pleistocene reefs is characteristic of this region, and is most likely related to rapidly fluctuating sea levels (see box 5.2). About 40% of the Nias Plio-Pleistocene fauna in table 5.4 are massive species by today's standards {sensu Moll 1983), compared to 28% in recent communities (Boekschoten et al. 1989). In contrast, 61% of the species in the Melolo Pleistocene collection are considered as massive. This difference suggests that sea-level fluctuations may have had a disproportionately greater impact on branching or sub-

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203

Figure 5.31. Aerial photograph of Guang Island illustrating the extensive reef complex that surrounds the limestone-capped island (on the right). Malimbo Island is located to the north (left). The reef system is exposed to weather during the Northwest Monsoon, but sheltered during the Southeast Monsoon. Sediment deposition occurs at the leeward side of the reef complex. Karst landscape of Guang Island is noticeable. Photo by Tomas and Anmarie

Tomascik.

massive shallow-water species than on the more massive species with greater depth range. However, the apparent low number of branching species may also be a sampling bias, since coralla of massive species are likely to survive longer. Additional data are clearly needed before more in-depth analyses can be conducted. The picture that emerged from the studies of Plio-Pleistocene fossil reef communities, was that following each glacial epoch and sea-level low stand, coral reef communities have always managed to reestablish themselves (Jackson 1992). Veron (1995) points out that only about 25%-30% of all extant hermatypic coral species have a fossil record, which may be indicative of species stability. Umbgrove (1946c) was the first to make an attempt to quantitatively interpret the Indonesian coral fossil record from an evolutionary standpoint. He came to the conclusion that during the past 30 million years the evolution of Indonesian coral fauna went through two distinct accelerations (i.e., speciations), one in the late Miocene and the second in the Pleistocene. To explain these two accelerated evolutionary tempos, Umbgrove (1946c) alluded to the diversification of marine habitats (i.e., intense geographic changes and formation of new deep-sea basins) as the principal

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factor which stimulated a n d facilitated Pleistocene coral speciation. However, h e also c o m m e n t e d o n t h e relatively sparse database o n which h e based his analysis. Probably the East Indies were at the time only slightly influenced by climatic changes, but the geographic changes were very intense. It is in the Pleistocene that mountain building (in the geographical sense) took place.... It seems probable to a high degree that we may look upon these events as the principal factor which stimulated the plastic group of corals to an intensive acceleration of their evolution; the more so, as the same epoch wrought great geographical changes over the whole of the tropical area of the Indo-Pacinc.—UMBGROVE 1946c T h e accelerated coral speciation in t h e I n d o n e s i a n Archipelago d u r i n g t h e Plio-Pleistocene, as suggested by Umbgrove (1946c), seems to have b e e n paralleled by a major acceleration of species t u r n o v e r in the C a r i b b e a n Province d u r i n g the middle Pliocene to early Pleistocene (4-3 Ma B.P.) (Budd et al. 1993). It seems that m o s t affected by the extinctions were g r o u p s of shallow-water species, especially from reef flat areas. U m b g r o v e ' s view of physically facilitated speciation is compatible with the latest c o n c e p t in coral evolution, which evokes surface circulation vicariance (a function of b o t h divergence a n d hybridization) as the primary driving force b e h i n d coral evolution (Veron 1995). T h e m a i n thesis of this hypothesis is

Table 5.3. Early Pliocene scleractinian genera from Guang and Bahuluang Islands, southwest Selayar. Numbers indicate the number of specimens found. Genera Acanthastrea Cyphastrea Favites Goniopora Pachyseris Pontes Alveopora Echinophyllia Galaxea Montastrea Symphyllia Acropora Montipora Favia Goniastrea Platygyra Pavona Seriatopora Blastomussa Fungia Leptastrea Oulophyllia Tubipora

Guang Island 2 3 8 5 2 10 2 1 2 1 1 12 1 8 16 6 2 1 2 1 2 1 6

Data from van der Land and Sukarno 1986.

Bahuluang Island 2 1 8

2

2

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CENOZOIC REEFS

that alterations in surface oceanic circulation patterns will control the amount of genetic connectivity among different interbreeding populations, which may ultimately result in genotypic divergence. The necessary mechanisms for the alteration of circulation patterns are implicit in Umbgrove's argument. It is accepted that climatic changes, as they relate to either atmospheric or seawater temperatures, had at the most a minimal effect on the evolution of Indonesian coral communities. However, the rapidly fluctuating sea levels during the Plio-Pleistocene and active orogenies have left their mark. On a global scale, paleoclimatic cycles are most likely of greater significance than geological processes (i.e., tectonics). However, in a relatively small area such as the Indonesian and Philippine Archipelagoes, intense geological activity (volcanism and tectonic uplift or subduction) combined with rapidly fluctuating sea levels can be relatively fast on geological and evolutionary time scales. The formation of island arcs and submarine ridges alters the surface circulation patterns, either through subsea deflection of deep currents or through direct interference at the surface. The Plio-Pleistocene was a period of renewed orogeny in the archipelago, creating new substrates and new habitats for coral colonization (Umbgrove 1946c). Rapid colonization of lava flows (i.e., 10 km 2 ) oceanic platform reefs (e.g., Lucipara Islands, Karang Skaro, Nil Desperandun, Karang Dusborgh, all in the Banda Sea), where they seem to contribute a significant amount of biogenic carbonate to the construction of extensive reef flats. However, there are no quantitative studies available to support this recent field observation. It is interesting to point out, however, that, unlike many other oceanic reef systems (e.g., Palau, Hawaii, etc.), the Banda Sea coral reef platforms are located in an area of intense seasonal upwelling (Wyrtki 1961), with correspondingly high productivity (Gieskes et al. 1990)', and more importantly, an abundant supply of inorganic nutrients which are pumped into the euphotic zone during each Southeast Monsoon (Wetsteyn et al. 1990). Along the coastlines of large islands fringed by reefs (e.g., Bali), the skeletal remains of benthic foraminifera are a significant source of beach sand, thus of considerable economic value to regional economies (fig. 9.1). Hallock (1976) documented earlier that approximately 0.2% of the coastal and nearshore sand reservoir of Oahu, Hawaii, was produced by the benthic symbiont-bearing foraminiferans Amphistegina, Heterostegina and Marginopora, with Amphistegina accounting for about 90% of the total production. Contribution of benthic coral-reef foraminifera to the total primary production of the reef system has not been fully determined, but their role as primary producers must be significant. In Palau, carbonate production estimates of rotaliinids on the high-energy seaward reef flats were up to 2.8 kg CaCO3 .m-2.yr1, while on the lagoonal slopes the productivity stood at 0.6 kg CaC03.m"2.yr~ (Hallock 1981b). These, and other early studies, clearly demonstrated the significant role of reef-associated benthic foraminifera in the production of reefal sediments (Chapman 1900). However, 371

372

FORAMINIFERIDA

Figure 9.1. A) The beach deposits along the Sanur and Nusa Dua coastlines, southern Bali, consist almost entirely of the remains of benthic foraminifera, Baculogypsina sphaerulata. B) B. sphaerulata is a spherical benthic foram that thrives at the seaward edge of the Sanur reef flat. Photos by Tomas and Anmarie

Tomascik.

their contribution to reefal build-ups has been generally unappreciated, even though the group had a significant input in reefal development since the late Paleozoic (Hallock 1981b). Extensive limestone deposits on Java, Madura, Sumatra, and the impressive Neogene limestones of Mangkalihat Peninsula in East Kalimantan, for example, are of benthic foraminiferal origins. In Bali, the magnificent white cliffs along the azure coast of southwest Bukit Badung, near the ancient Ulu Watu Temple, are made mostly from early Neogene foraminiferal and coral limestone deposits, and together with the magic of Ulu Watu Temple, offer tourists (and the intrepid scientist) an unforgettable experience (fig. 9.2). With regards to Recent reefs in Indonesia, it seems that corals, calcareous algae and benthic foraminifera are the major carbonate producers, a fact clearly recognized by Molengraaff (1928) and Wells (1957) in other coral reef systems, past and present. The role of benthic foraminifera in the construction of the coral reef framework was demonstrated by Finckh (1904) on Funafuti Atoll cores, in

INTRODUCTION

373

Figure 9.2. Massive, early Neogene, limestone cliffs along the Bukit Badung (Bali) southern coastline are made almost entirely from foraminiferal and coral deposits raised by tectonic uplift. Photo by Tomas and Anmarie

Tomascik.

which, in terms of reef-building importance, benthic foraminifera were outranked only by the coralline algae and Halimeda spp. Surprisingly, scleractinian corals were ranked fourth. Scoffin and Tudhope (1985) have since shown that on the outer reefal shelf of the Great Barrier Reef, Australia, benthic foraminifera are the dominant producers of calcareous reefal sediments. Similar conditions seem to exist on most offshore and oceanic reefs in the Indonesian Archipelago. Avast majority of extant foraminifera are benthic; however, some families are primarily pelagic. There are less than 100 extant pelagic foraminiferan species (Murray 1991b). Nevertheless, they are very abundant and constitute a significant biomass of the zooplankton community, as well as being the most important contributors of CaC0 3 to deep-sea sediments. About 30 planktonic foraminifera species are restricted mainly to the subtropical and tropical oceans (e.g., species of the Globorotalia, Globigerinella, Globigerinoides, Pulleniatina, Hastigerina); however, very little information is available on their biology, ecology and distribution in Indonesian waters. Most of the tropical pelagic foraminifera found in the surface waters (0100 m) are symbiotic with photoendosymbionts (e.g., dinoflagellates, mainly Gymnodinium beii, and Chrysophyta) (Spero 1987; Faber et al. 1988, 1989), and are therefore generally restricted to the euphotic zone (maximum depth about 150 m). A similar symbiotic association and distribution are found in the shallow-water benthic foraminifera, such as Calcarinidae, Asterigeridae, Alveolinidae, Amphisteginidae, Nummulitidae, and Soritidae (Hallock 1984). Most planktonic foraminifera live in the euphotic zone, and attain their maximum densities in the deep chlorophyll maximum layer (Murray 1991b), whose depth varies along both spatial a n d t e m p o r a l lines (Fairbanks and Wiebe 1980). Non-symbiotic foraminifera, pelagic and benthic, on the other hand, have a much wider depth distribution, since they do not require sunlight for growth and development.

374

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CLASSIFICATION In shallow tropical coastal-water environments where high temperatures predominate, and seawater is supersaturated with respect to CaC0 3 , skeletal remains of microscopic calcifying organisms accumulate in great abundance, and foraminifera tests abound. The foraminifera of the Indonesian Archipelago are a diverse group of pelagic and benthic (intertidal to abyssal) protozoans found in marine and estuarine environments. Even though the extant foraminifera are among the most abundant groups of organisms in many benthic communities, they have received relatively little attention in Indonesia, mainly because of taxonomic difficulty. The fossil record of foraminifera dates to the Cambrian, with over 35,000 species described. The exact number of extant species is not known, but is most likely somewhere between 3500-5000. Thus, foraminifera are one of the most widespread and long-lived groups of organisms in today's oceans. To obtain a perspective of the taxonomic complexity of this group, consider the 70-plus volumes of the Catalog of Foraminifera. The first significant work on the extant marine benthic foraminifera of the Indonesian Archipelago was that of Hofker (1951), who described 16 families, 52 genera and 104 species from the Siboga Expedition collections. Twenty-seven years later, the benthic samples from the 1929-1930 Snellius Expedition yielded an additional 462 species (Hofker 1978). However, the vast majority of the Siboga and Snellius specimens are from oceanic deep-water samples, thus coral-reef-associated foraminifera are under-represented, mainly as a result of sampling effort. This diverse group of protozoans is "currently" assigned to the Kingdom Protista (or Protocista) (table 9.1). However, the taxonomic certainty of many of these groups is questionable at best, and the final fate of the Subkingdom Protozoa remains to be decided. As of now, the Protozoa is: "...a large, unwieldy assemblage, so grouped only by consensus of protozoologists" (Leeetal. 1985). The foraminifera have characteristics such as cytoplasmic organization and pseudopodia that are very similar to most amoeboid organisms grouped in the Protozoa. One of the main characteristics that sets foraminifera apart from amoebae is the presence of hard biogenic tests, which most species secrete themselves, and from which the pseudopodia, or rhizopodia, protrude in a net-like array (Anderson and Lee 1991) (fig. 9.3). The extant foraminifera are currently grouped into four suborders and 13 superfamilies (table 9.2), but detailed classification of foraminifera is beyond the scope of this book, and the reader is referred to Ellis and Messina (1965, 1966) for an introduction to the foraminiferal systematics.

BIOLOGICAL ASPECTS

Test Morphology The classification of Foraminiferida is based mainly on test morphology (i.e., architecture and ornamentation), growth patterns and material composition (Lee et al. 1985). Compared to other protozoa, foraminifera are relatively large organisms

BIOLOGICAL ASPECTS

375

Table 9.1. Current taxonomic position of Order Foraminiferida. Taxonomic rank

Classification

Kingdom Subkingdom Phylum Subphylum Superclass Class Order

Protista Protozoa Sarcomastigophora Sarcodina Rhizopodea Granuloreticulosea Foraminiferida

Table 9.2. Order Foraminiferida. List of suborders and superfamilies with some basic morphological characteristics. Suborder/superfamily

Test characteristics

SUBORDER ALLOGROMIINA

Test smooth, flexible, membranous, or with agglutinated material. Same characteristics as the order. Test arenaceous, membrane with a rigid cover of cemented foreign material. Tests without internal subdivisions; deep-sea sediments. Test heavily chambered.

Superfamily Lagynacea SUBORDER TEXTULARIINA Superfamily Ammodiscacea Superfamily Lituolacea SUBORDER MILIOLINA Superfamily Miliolacea SUBORDER ROTALIINA Superfamily Nodosariacea Superfamily Buliminacea Superfamily Discorbacea Superfamily Spirillinacea Superfamily Rotaliacea Superfamily Superfamily Superfamily Superfamily

Globigerinacea Orbitodacea Cassidulinacea Carterinacea

Source: Lee et al. 1985.

Tests porcelaneous; dirty brown to amber in transmitted light. Tests resemble fired porcelain. Tests hyaline and glassy; numerous foramina. Test wall solid but may be laminate; pores < 1 urn in diameter; deep water. Test chambers trochospiral; pores < 1 urn; aperture with an internal flex. Test wall with laminated calcite that is perforate and noncaniculate; shallow water. Perforate calcareous test; tubular second chamber in planispiral arrangement. Tests very small trochospiral; wall calcareous; euryhaline. Tests with double walls of hyaline calcite; all pelagic. Tests with 2-layered walls of radially laminated calcite. Test walls of perforate, granular calcite. Test walls of fusiform calcareous spicules.

376

FORAMINIFERIDA

Figure 9.3. Photograph of a planktonic foraminifera Globigerinoides ruber, showing multi-chambered test with long delicate spines sheathed with cytoplasm. Photo courtesy of H.J. Spero, University of California,

Davis.

with adult test diameters ranging from < 0.1 mm (e.g., some species of planktonic foraminifera) to 100 mm (e.g., Marginopom vertebralis) (fig. 9.4). Forams may have life spans ranging from weeks to years (Lee et al. 1985; H. Spero, pers. comm. 1996). Foraminiferan tests are either simple unilocular (i.e., single chamber Allogromiina) to relatively complex multilocular (i.e., many chambers - Rotaliidae) forms. In multichambered species, test growth is always initiated from a proloculus (i.e., the first chamber) to which new locula are added, each connected with the other by a foramina (thus the name foraminifera), which is the previous aperture (fig. 9.5). The locula (i.e., chambers) are added in a specific sequence depending on the species (e.g., rectilinear or coiled). The locula can be added in a variety of ways (uniserial, biserial, or triserial), and can produce a diversity of shapes from spiral, flat coiled, conical, etc. In many species the surface of the test is covered by elaborate projections or ornamentations, which are used for taxonomic identification. The tests consist mainly of CaC0 3 , predominantiy calcite, which can be hyaline (i.e., transparent), porcelaneous, granular, fibrous or alveolar (Lee et al. 1985). The extant benthic foraminifera with characteristically mineralized calcareous tests are the Suborders Miliolina and Rotaliina (Lee et al. 1985). However, foram tests may also be gelatinous, chitinous, arenaceous (e.g., Textulariina) or siliceous. Superfamily Robertinacea are characterized by perforated tests made of aragonite (e.g., Mississippina, Robertina, Rubratella). In terms of surface features, foram tests can be either perforated by numerous small pores (i.e., perforations), or be imperforate with a distinct smooth porcela-

BIOLOGICAL ASPECTS

377 Figure 9.4. Marginopora vertebralis is a significant producer of reefal C a C 0 3 on many Indonesian reefs. Pulau Sago, Banggai Islands (depth 10m). Courtesy of Coral Cay Ltd.

Conservation

Figure 9.5. General patterns of test formation in Foraminiferida. After Jahn et al. 1979.

neous appearance (e.g., Qinqueloculina, Triloculina). In species without perforations, the pseudopodia stream out through a large aperture. In perforated groups, the small pores in the test allow the cytoplasm to stream out, thus forming the characteristic reticulopodia which appear to be in constant motion. The fine, filamentous pseudopodia (i.e., rhizopodia) which branch freely and anastomose with one another, exhibit an interesting phenomenon (a continual puzzle for the biophysicists), which is the bi-directional transport or "streaming" of small granules in the cytoplasm. It seems that the granules along the outer surface of the pseudopodia

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are always moving towards the centre (i.e., inwards), while the particles in the inner core of the pseudopodia are usually moving towards the distal (i.e., outer) portion of the pseudopodium. The sticky pseudopodia are highly efficient food traps. In symbiotic species, the photosymbionts contained in the cytoplasm may undergo diurnal migrations along the length of the pseudopodia. During daytime the photosymbionts are concentrated towards the outer periphery of the pseudopodia, while at nighttime they withdraw into the test. In some foraminifera, the photosymbionts are periodically "harvested" (i.e., digested) by the animal.

REPRODUCTION According to Lee et al. (1985), reproduction in foraminifera can be through gametogenesis, binary fission, budding, fragmentation or cytotomy (or plasmotomy where cytoplasmic division takes place independently of nuclear division in multinucleate protozoa). However, these modes of reproduction are restricted mainly to the smallest of the foraminifera (e.g., Allogromia spp.). Whatever the mode of reproduction, the initial growth of all foraminifera begins with the formation of the first chamber called the proloculus (fig. 9.6). In the unilocular species, the single chamber persists; however, in multilocular species new loculi are added in a speciesspecific sequence. The majority of foraminifera have relatively complex life cycles involving dimorphic alternation of generations (i.e., asexual and sexual phases). Indeed, one of the most distinctive features of foraminifera is their mode of reproduction. Many foram groups exhibit dimorphism (or trimorphism), consisting of a megalospheric haploid sexual phase (i.e, gametic), and a microspheric diploid asexual phase (nongametic) (fig. 9.7). In fact, the life cycle of the foraminifera is highly complex, with eight different life-cycle programs recognized (Lee et al. 1991). However, of the 38 species studied in detail thus far, about 75% exhibit a metagenic life-cycle programme, which involves regular alternation of haploid sexual and diploid asexual generations (Lee et al. 1991). Only about 25% of foraminifera studied thus far exhibit an apogamic life-cycle programme, which is characterized by the predominance of asexual reproduction, and where sexual reproduction is either suppressed or does not occur (Lett et al. 1991). A gamic lifecycle programme, where sexual reproduction predominates, has been conclusively demonstrated in all planktonic forams studied so far (e.g., Orbulina universa, Globigerinoides ruber, G. sacculifer, G. conglobatus, Globigerina bulloides, Hastigerina pelagica) (H. Spero, pers. comm.). The mode of sexual reproduction varies as well, with about 70% of studied species undergoing gametogamy, characterized by the fusion of haploid gametes from two or more gamonts with subsequent formation of zygotes (Lee et al. 1991). Only about 10% of sexual reproduction occurs through autogamy, which involves the fusion of two haploid nuclei from the same parental cell (Lee et al. 1991). The uninucleate haploid gamonts represent the sexual phase of the reproductive cycle in foraminifera. The maturation of the gamont results in the formation of flagellated gametes which are released into the sea where fusion with gametes from other gamonts takes place. Thus zygotes are produced externally, and

REPRODUCTION

379 Figure 9.6. General morphology of a juvenile foraminifera. Source: Jahn et al. 1979.

Figure 9.7. Generalized representation of the metagenic life-cycle programme with alternation of haploid sexual and diploid asexual generations in foraminifera. 1) uninuclear gamont or agamete; 2) juvenile gamont with megalospheric proloculus; 3) mature gamont with pregametogenic mitosis followed by pregametic nuclei; maturation of isogametes; 4) fusion of gametes from different gamonts which usually form "nuptial cysts"; 5) diploid zygote; 6) juvenile agamont with developed microspheric proloculus; 7) mature agamont with simultaneous meiotic divisions followed by cytokinesis and formation of gamonts by schizogony (gamogony). Based on Beck and Braithwaite

1968.

380

FORAMINIFERIDA

the release of gametes by the gamont is analogous to broadcast-spawning in corals and other reef invertebrates. Not much is actually known about the planktonic phase of the zygote; however, it is well-equipped to survive in the plankton. The maturing zygote forms a microspheric proloculus and the animal begins to undergo intracellular nuclear divisions, with successive additions of new chambers. The asexual phase of the life cycle is diploid. The multinucleate animal in the asexual phase is called an agamont. In holoplanktonic species agamont maturation is completed in depths of less than 100 m. Once the agamont matures meiosis occurs, thus resulting in the formation of mononuclear cells called agametes which upon maturation are released from the agamont. The maturing agamete forms a megalospheric proloculus around itself, and successive additions of new chambers and gamete maturation begin once again. This reproductive sequence is a highly simplified generalization of the life cycle of most foraminifera. According to Lee et al. (1991), only about 38 species have been studied in detail thus far. Note, however, that there has never been any direct observation of asexual reproduction in any of the planktonic foraminifera (H. Spero pers. comm.).

SYMBIOSIS Considering the extent and diversity of algal endosymbionts in foraminifera, our previous neglect of the phenomenon is difficult to put into perspective... Their abundance in today's seas and their contribution to CaC03 shell production is not generally appreciated.—LEE AND ANDERSON 1991 The foraminifera are hosts to an amazing variety of photoendosymbionts that includes dinoflagellates, chlorophytes, unicellular rhodophytes, diatoms and chrysophytes (Lee et al. 1985; Lee and Anderson 1991). One of the most interesting recent discoveries was that certain groups of temperate foraminifera, notably the Elphiidae, Rotaliellidae and Nonionidae, are able to retain and use (i.e., husband or farm) chloroplasts from some algae which they have partially consumed (Lee and Anderson 1991); however, this has so far not been reported in tropical species. Most of the tropical symbiotic foraminifera are larger pelagic and benthic species. Size is relative, however, and in the world of Protozoa, foraminifera are the giants. Coral reefs in particular provide a variety of warm, shallow- and deep-water habitats which are favourable for the growth of photosynthetic benthic foraminifera (table 9.3). * The three dominant groups of benthic symbiotic coral reef-associated foraminiferans are the Families of Soritidae, Alveolinidae and Calcarinidae. This interesting and diverse group of coral reef organisms has, however, attracted relatively little interest from the biological community in Indonesia, even though knowledge of their life history strategies and environmental requirements would be of considerable benefit to petroleum geologists. Studies on benthic and pelagic foraminifera in Indonesia have only covered the thanatocenoses, and most research on this group has been done by geologists. It is interesting to point out that in the special reports published by the Lembaga Ilmu Pengetahuan Indonesia (LIPI), all foraminifera papers are assigned to the geology section.

381

SYMBIOSIS

T h e Family Soritidae are a c o m m o n g r o u p of shallow-water b e n t h i c foraminifera abundant along many tropical coastlines as well as coral reefs islands and atolls. Of all the invertebrates, soritid foraminifera are hosts to the most diverse assemblage of photoendosymbionts. The photoendosymbionts belong to such diverse groups as dinoflagellates, unicellular chlorophytes and rhodophytes (Lee et al. 1985). Among the most important coral-reef-associated genera are Peneroplis, Marginopora and Sorites. Family Alveolinidae, with their distinct cigar-shaped tests are a common component of the coral reef community, with genus Borelis being host to diatom-like photoendosymbionts (Lee et al. 1985). Diatoms are characterized by the presence of siliceous frustules, which, however, do not form when the diatoms are endosymbiotic. It seems that the suppression of frustule formation while in the host's cytoplasm is associated with the endosymbiotic lifestyle, since when removed from the foraminifera the diatoms are able to form normal frustules again. Scientists have been successful in artificially culturing diatom-like endosymbiont isolates (Lee et al. 1979b) from a number of benthic reef-associated foraminifera (e.g., Heterostegina depressa, Amphistegina lessonii) (Lee et al. 1980b), which resulted in the formation of characteristic siliceous frustules, thus enabling identification of a number of new species of pennate diatoms (e.g., Fragilaria shiloi and Navicula reisii) (Lee 1980a). Since this early work considerable progress has been made and research has expanded into the Indo-Pacific. According to Lee et al. (1992), there are about 20 diatom endosymbiont species recognized thus far. In a study conducted in Palau, Lee et al. (1992) reported that the most abundant endosymbiont in the eight species of foraminifera studied was Nitzschia frustulum var. symbiotica, which was found in 24% of all samples collected. Other diatoms isolated from the Palau samples were Amphora erezii, A. roettgeri, Fragilaria shiloi, Cocconeis andersonii, Navicula sp., N hanseniana, Nitzschia laevis, N panduriformis var. con-

Table 9.3. General distribution of known algal symbiont-bearing foraminifera on coral reefs and their preferred habitat (based on abundance). Note that the species listed may be found in more than one habitat. Species Baculogypsina sphaerulata Calcarina calcar Calcarina spengleri Peneroplis pertusus Peneroplis proteus Amphistegina lobifera Peneroplis planatus Marginopora vertebralis Amphistegina lessonii Heterostegina depressa Sorites marginalis Operculina ammonoides Amphistegina radiata Calcarina hispida Source: Hallock

1984.

Reef flat X X X X X

Lagoon hard bottom

X X X

Reef slope

X X X X

Deep slope

X X

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FORAMINIFERIDA

tinua, and Achnanthes maceneryae. Considering the close proximity of Palau to Indonesia, comparative study would be of interest. The eight species of benthic foraminifera from Palau that contain endosymbiotic diatoms are Amphistegina lessonii, A. lobifera, Heterostegina depressa, Neorotalia calcar, Calcarina spengleri, C. defrancei, C. gaudichaudi, and Baculogypsina sphaerulata. The Palau study conclusively demonstrated that foraminiferan hosts may contain a number of the endosymbionts at the same time. For example, Calcarina gaudichaudi houses nine endosymbiont species. In general it seems that most foraminifera contain between five to eight endosymbiont species at one time (Lee et al. 1992). However, Amphistegina lessonii and Heterostegina depressa each contain only three endosymbiont species. In Amphistegina lessonii the dominant endosymbionts are Fragilaria shiloi and Nitzschia frustulum ver. symbiotica, while in Heterostegina depressa the two dominant endosymbionts are Cocconeis andersonii and Amphora roettgeri (Lee et al. 1992). There also appears to be a geographic variation in the endosymbiont complement in the host animal. For example, Heterostegina depressa from the western IndoPacific (i.e., Palau and Kudaka Island, Japan) does not contain Nitzschia frustulum var. symbiotica (sample N=32). As was pointed out by Lee et al. (1992), inclusion of earlier studies revealed that both Amphistegina lessonii and Heterostegina depressa contain more endosymbionts as the sampling area increases. When considering the entire collection, H depressa is host to 12 endosymbiont species (Lee et al. 1992), thus there is a significant geographic variability. The dominant group of benthic foraminifera in the Indo-Pacific are the rotaliines, especially Family Calcarinidae. Some of the most important endosymbiotic rotaliines are Amphistegina, Baculogypsina, Calcarina, and Heterostegina. Endosymbiotic diatoms are associated with both Amphistegina and Heterostegina. Families Cymbaloporidae (e.g., Cymbaloporetta) and Homotrematidae (e.g., Homotrema) are important coral reef-dwellers worldwide (Lee et al. 1985). In an excellent review of foraminiferan symbiosis, which the reader is urged to refer to, Lee and Anderson (1991) provide a list of all known foraminiferan-endosymbiont associations. In total, there have been 47 foram-endosymbiont groups studied, thus far (i.e., up to 1991). Of all the known symbiotic foram-endosymbiont associations about 38% involve diatoms (e.g., Nitzschia), 26% chrysophytes (not yet identified), 21% dinoflagellates (e.g., Gymnodinium), 9% unicellular chlorophytes (e.g., Chlamydomonas), and 6% unicellular rhodophytes (e.g., Porphyridium). What seems to be interesting from the coral reef perspective, is that the three main coral reef-associated groups, Alveolinidae, Amphisteginidae and Calcarinidae, harbour only Bacillariophyta (diatoms), while the abundant coral reef soritids (e.g., Marginopora) are symbiotic with dinoflagellates. The evolution of symbiosis with a variety of photoendosymbionts has resulted in some significant morphological and physiological adaptations in the host species. Lee and Anderson (1991), and others, have asserted that: "algal endosymbiosis was, in fact, a driving force in their [foraminifera] evolution". One of the most interesting adaptations is the suppression of sexual reproduction. According to Bermudes and Back (1991), the suppression of sexual reproduction maybe related to the fact that asexual reproduction guarantees direct transfer of symbionts to the next generation, while sexual reproduction requires that the zygote, or the maturing agamont, must be reinoculated with the endosymbionts from the ambient seawater. Whether some gametes contain endosymbionts when they are released

SYMBIOSIS

383

from the gamont is not known, and is an interesting field of research. It appears that in the case of foraminiferan symbiosis, each biont (i.e., partner) may live independently, thus the symbiosis is facultative (Bermudes and Back 1991). The suggestion that the relationship may be ecologically obligate (Bermudes and Back 1991) may be too wide a generalization, since there are extensive shallow-water tropical coastal areas where inorganic nutrient and food limitation is not a factor (e.g., near major rivers or in upwelling areas). In general, symbiotic foraminifera tend to be much larger than their heterotrophic cousins, as is true for other symbiotic organisms, such as corals (Hallock 1985; Lee et al. 1979a; Hallock et al. 1991). For example, tests of Marginopora vertebralis can be over 100 mm in diameter. Other species such as the spherical Baculogypsina can reach over 2 mm during reproduction. With increased test size (i.e., body size) the sexual maturation has been delayed, compared to the purely heterotrophic groups, and according to past studies, their life spans are correspondingly longer (i.e., a few months to a year). However, some large deep-water foraminifera do not contain endosymbionts, while some very small foraminifera do (Hallock et al. 1991). According to Hallock et al. (1991), photoendosymbiosis and growth to a large size in foraminifera are considered adaptations to environmental conditions where food and energy resources are consistently limited. In comparison to the heterotrophic species, tests of benthic symbiotic foraminifera are significantly thinner and flatter, an adaptation that allows greater transmission of light, and exposes greater surface area to the downwelling solar radiation. In contrast to the flattened test morphologies of benthic symbiotic foraminifera, planktonic symbiotic foraminifera have evolved thin, spherical tests with numerous highly perforated chambers (e.g., Globigerina bulloides, G. falconensis, Globigerinella calida, Globigerinoides sacculifer, G. fistulosus, Globorotalia menardii), and elaborate ornamentation (e.g., spines) along which cytoplasm may stream out. In many planktonic species the endosymbionts are "farmed", and move along the pseudopodia in response to daily light cycles (Be et al. 1977). In all symbiotic foraminifera the endosymbionts are housed in specialized vacuoles within the cytoplasm (i.e., they are intracellular) (Bermudes and Back 1991). According to Lee (1983), the symbionts can constitute a significant percentage of the holobiont biomass. The above morphological adaptations seem to be directed at maximizing the photosynthetic efficiency of the photoendosymbionts as well as increasing the capacity for absorption of dissolved nutrients from ambient seawater. What nutritional role do the endosymbionts play in the association? Lee and Anderson (1991) point out that endosymbiosis in the larger foraminifera, with its tight recycling of nutrients and organic compounds, was an adaptation to oligotrophic tropical and subtropical oceans where they tend to predominate. It seems that the foramendosymbiont system is functionally similar to the coral-zooxanthellae symbiosis; however, much less research has been done in this field (Lee et al. 1979; Hallock 1981a; McEnery and Lee 1981). Nonetheless, it seem clear that the foraminifera may potentially obtain a significant amount of energy from the endosymbionts in the form of various photosynthates that are translocated to the host animal, as is the case with the coral-zooxanthellae symbiosis. The potential amount of energy obtained by the animal biont is also dependent on the nature of the photoendosymbiont. For example, Hallock et al. (1991) pointed out that diatom-foram

384

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holobionts have higher rates of C 0 2 fixation (i.e., productivity) than do rhodophyte-foraminifera holobionts. What little information exists, suggests that in some associations (e.g., diatom-foram) the main photosynthates translocated from the endosymbiont to the animal biont are glucose and glycerol, which can be utilized by the animal biont in supplementing metabolic energy costs and/or growth (Keremer et al. 1980). However, some foraminifera, such as Heterostegina depressa, are apparently able to absorb sufficient nutrients from seawater so that they do not have to actively feed. Calcification In addition to the potential energy and food subsidy offered by the endosymbionts, the endosymbionts most likely have a considerable influence on the calcification process. Recent studies indicate that this may vary considerably among different species. For example, Ter Kuile and Erez (1987) demonstrated, using radionuclide C pulse-chase experiments, that Amphistegina lobifera was able to incorporate significant amounts of photosynthetically derived G into its skeleton, while a similar species was not. There have been a number of theories suggested in the past to explain the calcification process in animal-algal symbionts, and one of the earliest is that light stimulates calcification through C 0 2 fixation (i.e., photosynthesis) . In this theory, the driving mechanism behind light-enhanced calcification in photoendosymbiont-bearing foraminifera is an increase in pH as a result of photosynthetic uptake of C0 2 , which increases carbonate concentrations in the animal tissue near sites of carbonate deposition. However, Ter Kuile (1991) points out that new evidence indicates that calcification in some algal symbiont-bearing foraminifera is not stimulated by the photosynthetic fixation of C0 2 . Nonetheless, a number of studies have shown that calcification rates of endosymbiotic foraminifera incubated in light were up to three times higher than those incubated in the dark (Lee and Anderson 1991; Lea et al. 1995). Another mechanism evoked is the organic matrix theory which suggests that an organic matrix initiates or inhibits calcification by spatially arranging the calcium and carbonate ions (Ter Kuile 1991). Recent experiments with Orbulina universa suggest that Ca2+ in this species is in direct isotopic equilibrium with surrounding seawater, and that less than 0.5% of shell Ca2+ can be attributed to another source (Lea et al. 1995). According to these authors, "an internal cytoplasmic Ca2+ pool probably does not provide a means by which seawater cation ratios can be fractionated". Furthermore, their experiments demonstrated that calcification rates in adult O. universa are 2-3 times higher under high light intensities than in the dark. Since the enhancement in calcification rate is close to the calculated increase in carbon ion due to the photosynthetic activity of the endosymbionts, Lea et al. (1995) pointed out that calcification rate in O. universa may be proportional to the degree of carbonate saturation. Other possible mechanisms may involve active concentration of the reactants, since a number of foraminifera have been shown to concentrate calcium in internal pools (Ter Kuile 1991). One of the most plausible benefits of the symbiotic association for the foraminifers is the removal of metabolic wastes by the endosymbionts during photosynthesis. The removal of excess ammonium, phosphate and magnesium from the site of calcification will enhance the calcification process, since these substances are well-known crystal poisons. Whatever the final mechanism, the

HETEROTROPHIC-FEEDING

385

field has attracted considerable interest worldwide; however, up to now no comparable research has been initiated in Indonesia.

HETEROTROPHIC-FEEDING Most pelagic and benthic foraminifera are heterotrophic organisms feeding on a variety of organic matter (i.e., carnivorous, herbivorous, detrivorous or omnivorous), which is available in the water column, or in the sediments. Planktonic foraminifera are generally omnivorous (Anderson et al. 1979; Be 1982), feeding on any organic material that may get ensnared in their sticky pseudopodial nets (fig. 9.8). They can be considered as passive suspension-feeders, since they rely on collisions with food particles such as zooplankton, phytoplankton or suspended organic particulate matter (Murray 1991b). Feeding is essential for the growth of planktonic foraminifera; however, they may survive without food for long periods provided photosynthesis of their endosymbionts is not inhibited (Caron et al. 1981). Pelagic foraminifera are particularly well adapted for capturing zooplankton. The pseudopodial net of Globigerinoides sacculifer is highly effective in ensnaring small calanoid copepods (Spindler et al. 1984), which are its primary food source. The surface of their thin, perforated and multilocular tests is usually ornamented with long, thin and very delicate spines (often damaged in specimens from plankton tows), along which sticky cytoplasm can easily be seen streaming. In many species (e.g., Globigerinoides sacculifer, Orbulina universa) the pseudopodial nets supported by the spines substantially increase the surface area of the feeding surface. The spines can be very long, in as much as it is actually possible for scientists to drift in the water, and capture individual foraminifera (e.g., Globigerinoides spp.) in small glass jars as they drift by (Tomascik and Mah, pers. obs.). Capturing planktonic foraminifera in jars is a morning occupation of many foram scientists involved in experimental research, since great care must be taken so that the long delicate spines are not damaged, which would otherwise ruin the experiments. Spinous planktonic foraminifera are especially abundant in oligotrophic water masses, where calanoid copepods are the dominant zooplankton. According to Spindler et al. (1984), planktonic foraminifera on an average capture and consume one zooplankter per day (fig. 9.9). During the Northwest Monsoon (i.e., nonupwelling period) the most abundant planktonic foraminifera in the central Banda Sea are Globigerinoides ruber and G. sacculifer, which are characteristic of oligotrophic water masses in other oceanic regions (e.g., Red Sea, central Indian Ocean) (Troelstra and Kroon 1989). During the same period the zooplankton community is dominated by copepods, which represent about 77% of zooplankton abundance (Baars et al. 1990). The oligotrophic conditions during the Northwest Monsoon are reflected by lower zooplankton biomass (9.5±2.8 cc.m") compared to the Southeast Monsoon upwelling season when the average zooplankton biomass in the eastern Banda Sea doubles (20.0±8.4 cc.m"2) (Schalk 1987). In more eutrophic regions, the planktonic foraminifera are dominated by non-spinose species (e.g., Neogloboquadrina dutertrei, Globorotalia menardii, Pulleniatina obliquiloculata), which are also mostly omnivorous. However, some species such as N. dutertrei feed exclusively on phyto-

386

FORAMINIFERIDA

Figure 9.8. Photograph of a symbiotic planktonic foraminifera, Globigerinoides ruber, with ensnared zoo-plankton within the net-like pseudopodia. Photo courtesy of H.J. Spero, University of California,

Figure 9.9. Photograph of a symbiotic planktonic foraminifera, Orbulina universa, feeding on a captured calanoid copepod. Photo courtesy of H.J. Spero, University of California, Davis.

Davis.

HETEROTROPHIC-FEEDING

387

Figure 9.10. Benthic foraminifera are among the most abundant reef organisms. A) Micrograph of Baculogypsina sphaerulata (Parker and Jones), a dominant reef-associated benthic foraminifera at Sanur, Bali. B) Schlumbergerella floresiana (Schlumberger) at Nusa Dua, Bali. Photos courtesy of M. K. Adisaputra,

Pusat Pengembangan

Geologi Kelautan - LIPI,

Bandung.

plankton (Hemleben et al. 1980), and are usually associated with the Deep Chlorophyll Maximum layer where phytoplankton occurs in the greatest concentrations (Barmawidjaja 1993). Benthic foraminifera exhibit a variety of feeding strategies. Detritus-feeders (i.e., detritus plus associated bacteria) are present at all depths. Feeding occurs at the sediment/water interface or infaunally. Test morphologies in benthic foraminifera reflect feeding strategies. More erect types may be more common in environments where pseudopodial nets extended into the water column provide an efficient mechanism to trap suspended organic matter "raining" down, or drifting by in slow currents (Murray 1991a). Active herbivory in benthic habitats is restricted to the euphotic zone, where benthic foraminifera are known to forage (Travis and Bowser 1991). For example, Allogromia laticollarisi uses an extensive network of sticky reticulopods (i.e., highly branched and anastomosed fllopods) that are used for locomotion as well as in accumulating benthic diatoms in net-like enclosures (Travis and Boswer 1991). Uptake of dissolved organic material may also be an important feeding strategy (Schwab and Hofer 1979). Many of the shallow-water foraminifera are symbiotic, thus translocation of photosynthates may be a significant supplement to their heterotrophic diets (Lee and Bock 1976). Many symbiotic benthic foraminifera in shallow-water coral reef environments are discoid or compressed, presumably to maximize photosynthetic efficiency of the holobiont. However, many highly abundant shallow-water symbiotic species, such as Baculogypsina sphaerulata or Schlumbergerellafloresiana,are spherical, with a prominent spinous architecture, which provides an efficient mechanism for attachment to benthic algae in high-energy environments (seaward reef flat) (fig. 9.10). Whether they are pelagic or benthic, feeding in most foraminifera is accomplished by the extrusion of pseudopodia (i.e., rhyzopodial nets) from the test of the

388

FORAMINIFERIDA

animal. The spines serve as an efficient net in which zooplankton invariably are ensnared, and subsequently digested. The food particles (or prey) are engulfed by the cytoplasm and digestion usually takes place either outside or inside the tests with the assistance of digestive lysosomes. T h e particle size captured by foraminifera, benthic and planktonic, ranges between 1- 20 um in diameter.

B E N T H I C FORAMINIFERA If there is one region of the planet that should have a high diversity of foraminifera, it is undoubtedly the Indonesian Archipelago. No other region in the world offers such a diversity of pelagic and benthic habitats as the thousands of islands and the open and enclosed seas of the archipelago. The pelagic habitats range from the warm oligotrophic surface waters to the cooler epipelagic nutrient-rich depths and the cold mesopelagic zone. In addition, vast areas of the Banda and Flores Seas are influenced by upwelling which results in highly productive surface waters during much of the year. In terms of benthic habitats, the Indonesian Archipelago offers a great diversity of shallow neritic environments (e.g., estuaries, mangroves, seagrass beds, coral reefs, etc.) as well as shallow-water oceanic habitats associated with coral islands, barrier reefs and atolls. Within the relatively small area of the Banda Sea, benthic foraminifera could be sampled from the sunlight intertidal reef flats of Pulau Manuk, and the dark hadal zone of the Weber Deep. General Distribution Among the best-known groups of foraminifera are the deep benthic forms, mainly as a result of sampling efforts during the past major expeditions. There are no comparable data on reef-associated foraminifera from the archipelago; however, benthic foraminifera from Jakarta Bay, Ambon Bay, Batam, etc., have been studied (Hamidjojo et al. 1980; Ongkosongo et al. 1980; Siregar and Hadiwisastra 1980; Helfinalis et al. 1989; Hermanto and Suhartati 1989; Suhartati and Subardi 1990). The available data indicate that some reefal foraminifera are represented in the sediments collected from bays fringed by coral reefs. Hofker (1978) was one of the first to compare the distribution of the benthic foraminifera collected during the Snellius Expedition. He pointed out the great dissimilarity in benthic foraminiferan fauna from the western and eastern regions of the archipelago (table 9.4). According to the samples collected during the Snellius Expedition, Hofker (1978) found that Jakarta Bay (western region) had only five benthic foraminifera common with samples from Sorong, Iran Java. The question is whether these differences are real (i.e., of ecological significance), or a sampling artifact. Furthermore, it is not possible to ascertain whether the sampling was done in comparable environments. Helfinalis and Rositasari (1988) investigated the foraminiferan fauna of Pulau Pari, an offshore shelf platform reef with an extensive shallow-water reef flat and two distinct lagoons, one of which has a depth greater than 15 m. According to the classification of Hopley (1982), the reef complex of P. Pari may be viewed as a mature lagoonal reef. The reef supports a diverse foraminiferan fauna, which is dis-

BALI CASE STUDY

389

tinctly different from the benthic foraminiferal assemblage of Jakarta Bay (table 9.5). However, because of different sampling techniques as well as different sampling efforts, the data in table 9.5 are not statistically comparable. Nonetheless, the data suggest that the differences in foraminiferan diversity among the various regions of the archipelago are most likely related to differences in environmental conditions of stations sampled. For example, the sediments of both Segara Anakan and Jepara are dominated by euryhaline foraminiferans, since both areas are influenced by high fresh water runoff. Coastal waters in these areas are estuarine (i.e., low salinities) during much of the West Monsoon (i.e., rainy season). To determine whether there are biogeographical patterns in the distribution of benthic foraminiferans, comparative studies are required. Coral reef-associated foraminifera offer a good opportunity for such studies, yet none have been done thus far.

ECONOMIC VALUE OF FORAMINIFERA Foraminiferans have proved to be indispensable tools for paleoclimatic, paleoecological and paleoceanographic reconstructions. Van Bemmelen (1949) ranked Foraminiferida as the most important stratigraphic tools of the Tertiary in the Indonesian Archipelago. The fossil remains of foraminifera (and other calcareous and siliceous plankton) are important indicators in biostratigraphy and dating of marine sediments, and have assisted petroleum geologists in their search for new oil deposits. Foraminifera, and the limestone sediments which they produce in conjunction with other calcareous planktonic and benthic organisms, are also of considerable economic value, and have been so, since early human cultures developed around the world. Among the best-known marine limestone deposits in the western world are the famous white cliffs (actually dominated by coccoliths) of Dover in England, which are rivaled by the anthropogenically orchestrated limestone deposits along the Nile, namely the Egyptian pyramids. The main source of building blocks for the pyramids were Eocene nummulitic limestones, consisting mainly of the foram genus Nummulites. In Indonesia, mining of limestone is of considerable economic value, and many of the shiny marble floors that are currently highly popular, are an intricate mosaic of foraminiferal fossils from the distant past. For example, the fine marbles from Gunung Panggul (south coast of East Java; Panggul Bay) originate from Miocene limestone deposits consisting mainly of Foraminifera (e.g., Aleveolinella bontangensis) and corals (van Bemmelen 1949).

B A L I CASE STUDY The role of extant foraminifera in the development of beach deposits is of significant economic value, a fact that has unfortunately not been recognized in the past, or the present (but see Adisaputra 1991). One of the greatest attractions of tropical countries with coral reefs are the long, white, palm-lined beaches, which in combination with the sun, surf and the azure- to-turquoise-coloured clear waters are an irresistible magnet for millions of tourists from the temperate regions. Indonesia

390

FORAMINIFERIDA

Table 9.4. Regional comparison of shallow-water Foraminiferida between Jakarta Bay (southwest Java Sea) and Sorong (Halmahera Sea). Jakarta Bay (west) Amphistegina radiata Asterorotalia pulchella Baculogypsinoides spinosus Bdelloidina aggregata Biarritzina proteiformis Bigenerina nodosaria Carpenteria utricularis Calcarina spengleri Clavulina pacifica Dendritina striatopunctata Discogypsina vesicularis Elphidium batavum Flintina bradyi Gypsina plana Heterostegina curva Miniacina miniacea Nummulites complanatus Parrellinum hispidulum Placopsilina bradyi Planorbulinella larvata Pseudorotalia schroeteriana Quinqueloculina kerimbatika Quinqueloculina bicarinata Quinqueloculina curta Quinqueloculina bidentata Reophax scorpiurus Rotalidium concinnum Sagenina frondescens Schlumbergerina areniphora Sphaerogypsina globularis Spiroloculina communis Textularia kerimbaensis Triloculina tricarinata Triloculina rupertiana Valvotextularia foliacea Valvotextularia (T.) rugulosa Data from Snellius Expedition

1929-1930.

Sorong Alveolinella quoyi Amphistegina madagascarensis Aomalinella rostrata Amphistegina lessonii Asterorotalia annectens Cellanthus craticulatus Clavulina difformis Clavulina pacifica Cymbaloporetta bradyi Elphidium crispum Eponides repandus Heterostegina depressa Heterostegina heterosteginoides Miniacina miniacea Nubecularia divericata Neoeponides broeckiana Nummulites complanatus Nummulites venosus Orbitolites variablis Pararotalia defrancei Parrellinum hispidulum Peneroplis planatus Pyrgo denticulata Quinqueloculina lamarckiana Quinqueloculina quadrilateralis Quinqueloculina parkeri Quinqueloculina granulocostata Schlumbergerina areniphora Sigmoidella elegantissima Spiroloculina corrugata Triloculina tricarinata Textularia barker! Textulariella rugulosa

391

BALI CASE STUDY

Table 9.5. General distribution of some benthic foraminifera in the Indonesian Archipelago. 1) Jakarta Bay 1929-1930. Habitat is coral and mud; 2) Sorong, west Irian Jaya 1929-1930. Habitat is mud and sand; 3) Jakarta Bay 1975-1979. Habitat is coral and mud; 4) Batam, northwest Java Sea, Teluk Tering. Habitat sand; 5) Ambon Bay, Ambon, Moluccas, 1984-1987; 6) Segara Anakan, south Central Java. Habitat sand, mud, sandy mud; 7) Jepara, north Central Java. Habitat sand, sandy mud; 8) Pulau Pari, north West Java. Habitats - 8A: Reef flat, 8B: Lagoon; 8C: Reef slope. FORAMINIFERA Adelosina semistriata Amphistegina lessonii Amphistegina madagascarensis Alveolinella quoyi Ammobaculites agglutinatus Ammonia beccarii Ammonia concinna Ammonia gaimardi Ammonia umbonata Amphistegina lessonii Amphistegina quoyii Amphistegina radiata Aomalinelia rostrata Archais ungulatus Asterigerina carinata Asterorotalia annectens Asterorotalia pulchella Asterorotalia trispinoa Baculogypsina sphaerulata Baculogypsinoides spinosus Bdelloidina aggregata Biarritzina proteiformis Bigenerina nodosaria Bolivina earlandi Bolivina schwagerina Calcarina calcar Calcarina spengleri Calcarina venusta Cancris oblongus Carpenteria utricularis Cellanthus craticulatus Cibicides lobatulus Cibicides praencinctus Clavulina difformis Clavulina pacifica Cymbaloporetta bradyi Cymbaloporetta squamosa Dendritina striatopunctata Discogypsina vesicularis Discorbina mira Elphidium advenum Elphidium batavum Elphidium craticulatum Elphidium crispum Elphidium depressulum Elphidium hispidulum

1

2

3

4

X X

X

X

X

X X X X

X

X

X

X X

7

8 A

B

X

X X X

8

C

X

X X X X X

X

X

X

X

X X

X X

X

X

X X X X X X X

X

X

X X

X

X

X

X X

X

X

X X X X

8

X

X

X X X X

X

6

X

X X X

X

5

X

X

X X

X X X X X X

X

X

X

X

X X X X X X X X X X X X X X X X X X X X X X X X X

392

FORAMINIFERIDA

Table 9.5. (Continued.) FORAMINIFERA Elphidium lessonii Elphidium macellum Elphidium microgranulosum Eponides repandus Eponides umbonatus Flintina bradayana Gypsina plana Harplophragmoides canariensis Hauerina bradyi Hauerina fragillisima Hauerina involuta Hauerina speciosa Heterostegina curva Heterostegina depressa Heterostegina heterosteginoides Hoglundina elegans Lecticulina convergens Lecticulina cultrata Loxostomum amygdalaeformis Loxostomum limbatum Massilina tropicalis Marginopora vertebralis Massilina milleti Miliolinella oblonga Miliolinella subrotunda Miliolinella sublineata Miniacina miniacea Neoconorbina terouemi Neoeponides broeckiana Noniella atrizans Nonion boueanum Nonion depressulum Nubecularia divericata Nummulites complanatus Nummulites granulans Nummulites venosus Oolina apiculata Operculina ammonoides Orbitolites variablis Pararotalia defrancei Parrellinum hispidulum Peneroplis pertusus Peneroplis planatus Pileolina pattelliformis Pileolina tubernacularis Placopsilina bradyi Planorbulinella larvata Planularia australis Planulina bradii

1

2

3

X

4

X

X X

5

6

7

8

8 8 A

X X

X

X X

X X

X

X X

X X

X X X

X X

X X X

X X X

X X X X X

X X X X

X X

X

X X X X X X X

X

X X X X X

X X

X

X

X X X

X X

X

X X

X

X

X X

B C X X

X X

X

X X

X

X

X X X X

X X

X X X

X X X

X

393

BALI CASE S T U D Y

Table 9.5. (Continued.) FORAMINIFERA Poroeponides cribrorepandus Praesorites orbitolitoides Pseudomassilina macilenta Pseudorotalia schroeteriana Pyrgo denticulata Pyrgo depressa Pyrgo lucernula Pyrulina angusta Quinqueloculina auberiana Quinqueloculina bicarinata Quinqueloculina bidentata Quinqueloculina boueana Quinqueloculina bradyana Quinqueloculina cultrata Quinqueloculina curta Quinqueloculina granulocostata Quinqueloculina intricata Quinqueloculina kerimbatika Quinqueloculina lamarckiana Quinqueloculina linneana Quinqueloculina parkeri Quinqueloculina pittensis Quinqueloculina poeyana Quinqueloculina pseudoreticulata Quinqueloculina pulchella Quinqueloculina quadrilateralis Quinqueloculina reticulata Quinqueloculina seminulum Quinqueloculina tasmanica Quinqueloculina tropicalis Quinqueloculina venusta Reophax scorpiurus Reussella simlex Robulus thalmanni Robulus vortex Rotalia conoidea Rotalia gaimardii Rotalia trispinosa Rotalidium concinnum Sagenina frondescens Schlumbergerina areniphora Sigmoidella elegantissima • Sigmoilopsis schlumbergeri Siphogenerina raphanus Sphaerogypsina globularis Spiroloculina angulata Spiroloculina arenaria Spirolina arietina Spiroloculina communis

1

2

3

4 X X

X

X

X

X X X X

X X X

X X X X

X X

X X X

X

X X

X

8

8 8

X X X

X

X X

B

X X

X

X X

X X

X X

X X

X

X

X X

X

X

X

X

X X X X

X

X

X X X

X

X

X X

X X X X X X

X X

X

X X

X

X X X

X

C

X

X X

X X X X

7

A

X X X

X

6

X X X

X X

X

5

X

X X

X X

X

X

X

X

394

FORAMINIFERIDA

Table 9.5. (Continued.) FORAMINIFERA Spiroloculina corrugata Spirosigmoilina parri Textularia angglutinans Textularia barken Textularia kerimbaensis Textularia pseudogramen Textulariella rugulosa Triloculina cuneata Triloculina rupertiana Triloculina transversestriata Triloculina tricarinata Triloculina trigonula Valvotextularia foliacea Valvotextularia (T.) rugulosa Vulvulineria rugosa Planktonic Foraminiferida Globigerina bulloides Globigerina falconensis Globigerinella callida Globigerinoides conglbatus Globigerinoides cyclostomus Globigerinoides fistulosus Globigerinoides ruber Globigerinoides sacculifer Globigerinoides sacculifer Globoquadrina pseudofoliata (Globorotalia berudezi Globorotalia hersuta Globorotalia menardii Globorotalia neoflexuosa Globorotalia pseudopumilio Globorotalia puncticulata Globorotalia seiglei Globorotalia trancatulinoides ' Globorotalia tumida Globorotalia ungulata Glonulina rotundata Globigerinella siphonifera Neogloboquadrina blowi Neogloboquadrina hymerosa Neogloboquadrina incompata Orbulina universa Pulleniatina finalis Pulleniatina obliqueloculata Pulleniatina praecursor Pulleniatina primal is Sphaeroidinella dehiscens

1

2

3

4

5

6

7

8

X

X

X

X X

X X

X X

X

X

X

X X X X X X X

X

X X

X X

X X

8 8 C

B

X

X

X

A

X

X

X

X

X X

X X X X X X X X X X X X X X X

X X X X

X X X X X X X X

X

X X '

X X X X X X

X X X X

X

Source: Jakarta Bay 1929-1930 and Sorong 1929-1930 Hofker 1978; Batam Suhartati and Subardi 1990; Jakarta Bay 1975-1979 Hamidjojo et al. 1980; Ambon Bay Hermanto and Suhartati 1989; Segara Anakan Subardi et al. 1989; Jepara Helfinalis et al. 1989.

SOURCE OF SANUR BEACH SAND

395

is no exception, and Bali in particular has recently experienced a tourism boom. Putting aside the fascinating cultural experience that one obtains from visiting Bali, much of the future success of the tourist industry will depend on maintaining the island's environmental integrity, especially of the coastal marine environments. One of the most pressing management tasks will be the maintenance of the beaches. Sanur's golden beach, which is one of the most valued attractions on the island, is composed almost entirely of bioclastics (i.e., sediments made entirely of broken fragments of biogenic skeletal material). Indeed, about 80% to 90% of Sanur beach sand consists of foraminiferal tests. The foraminiferal tests of various sizes are dominated almost entirely by the calcarinids Baculogypsina sphaerulata and Schlumbergerella floresiana, which contain photoendosymbiotic diatoms. As a result, they prefer shallow-water, high-energy coral reef habitats.

S O U R C E OF S A N U R B E A C H S A N D Where does the immense amount of biogenic carbonate come from? The answer to this question lies in an extensive coral reef complex that runs along the entire length of the Sanur and Nusa Dua coastlines. Coral reefs that fringe large stretches of Bali's coastline are the classical fringing reefs. Along the south coast, especially Sanur beach and Bukit Badung Peninsula (e.g., Nusa Dua), the fringing reef extends a considerable distance offshore (e.g., up to 700 m at Cemara beach). Because of its relatively long and unbreached length (i.e., 8 km long), as well as a wide boat channel (sensu Guilcher 1988) (i.e., lagoon or moat), the Sanur fringing reef complex has often been incorrectly referred to as a barrier reef. The seaward edge of the fringing reef is above the mean high-water neap mark, and is delineated by white breakers. The depth of the boat channel varies; for example, between Bali Hyatt beach and Cemara beach the bottom of the boat channel is roughly at the mean low-water spring mark, whereas just off the Bali Beach Hotel to the north, the bottom is at about mean low-water neap mark. Note that the tidal range along the south coast of Bali is about 2.6 m. Because of a relatively restricted connection with the sea, water level inside the boat channel during low tides is above the offshore sea level, and thus most of the boat channel is always flooded. Baculogypsina sphaerulata and Schlumbergerella floresiana are large spherical foraminifera that flourish at the seaward edge of the reef flat, and especially at the outer reef crest. These zones are classified as high-energy environments, receiving the full force of the Indian Ocean swell, a fact much appreciated by the world surfing community. At the seaward edge of the reef (i.e., algal rim), both species live attached to benthic macrophytes (e.g., Eucheuma spp., Galaxaura spp., Codium spp., Gracilaria spp., Acanthophora spp., Rhodymenia spp., Lurencia spp., Tricleocarpa spp., Hypnea spp., Dyctiosphaeria spp., and Sargassum spp.). The calcareous rhodophytes, particularly Lithophyllum moluccense, Jania spp. and Amphiroa spp., also provide suitable substrate for the a t t a c h m e n t of these two b e n t h i c foraminiferans. In addition, both Sanur and Nusa Dua boat channels are important seagrass habitats (i.e., eight seagrass species recorded: Enhalus acoroides, Cymodocea serrulata, Thalassia hemprichii, Thalassodendron ciliatum, Halophila ovalis, H. minor, Halodulepinifolia, H. uninervis) whose calm 'lagoonal' waters provide an ideal envi-

396

FORAMINIFERIDA

ronment for a variety of other benthic foraminiferans. Among the most abundant benthic foraminifera in the seagrass habitats are Amphistegina lessonii (Amphisteginidae), Calcarina calcar (Calcarinidae) and Heterostegina sp. (Nummulitidae), which are symbiotic with Nitzschia frustulum var. symbiotica (diatom) (Adisaputra 1991; Lee and Anderson 1991). Other smaller shallow-water benthic foraminifera found along the southern coastline of Bali are Elphidium advenum (Elphidiidae), Quinqueloculina pseudoreticulata (Miliolidae), Q. seminulina, Q. parked, Spiroloculina sp. (Nubeculariidae), Textularia sp. (Textulariidae) and Cellanthus craticulatus (Adisaputra 1991). Both Textularia sp. and Elphidium advenum are euryhaline species capable of withstanding very low salinities. Greater abundance of these two species in the sediments along the southern section of Sanur beach is most likely related to the brackish water coastal habitats that dominate this region of the coastline. Nonetheless, their contribution to beach sediments remains much lower than B. sphaerulata and S. floresiana. The composition of benthic foraminifera on the seaward slope of Sanur, Benoa and Nusa Dua reefs, is much different from the lagoonal habitats. According to Adisaputra (1991), offshore areas are dominated by foraminifera belonging to Orders Rotaliina, Miliolina and Textulariina. At depths of less than about 12 m, Cribrononion hispidulus (Elphidiidae) seems to be the dominant benthic foram (Adisaputra 1991), especially along the Benoa coastline, which is characterized by low salinities. Further offshore, species numbers increase with depth, and genera such as Triloculina, Spiroloculina, Ammonia, Articulina, Anomalina, Clavulina, Eponides, Brizalina, Miliolinella, Martinotiella, etc., become more abundant (Adisaputra 1991). Pelagic foraminifera do not contribute much to the beach sediments; however, they are present in offshore sediment samples at depths of more than 16 m. The most abundant pelagics are Globigerina bulloides, Globigerinoides cyclostomus, G. ruber, G. sacculifer and Neogloboquadrina dutertrei. Note that G. bulloides is an indicator of upwelling conditions, which occur along the south coast of Bali. Much has been said about the degradation of the Bali reefs; however, rapid surveys of the Sanur and Nusa Dua fringing reefs revealed that the reefs are still in relatively good condition, and that with appropriate management of coastal activities the reefs will be able to survive, if not regenerate. Note that the relatively good 'health' of these reefs is mainly due to strong currents that continually flush the coastal areas and bring in clean oceanic waters. However, there are signs that major changes have occurred. The seagrasses in the lagoon have increased, and while this may not be a problem at this time, continual discharge of nutrient-rich waste-water (e.g., sewage) will eventually have an impact, as it did in many other areas. However, of greater concern to the beach stability is the ill-advised building of structures for beach protection such as groins and cement walls. Both Sanur and Nusa Dua beaches rely on the longshore transport system to continually replenish the beaches with dead tests (i.e., shells) of benthic foraminifera, which originate at the seaward edge of the reef crest. Interference with the longshore transport system will ultimately result in increased beach erosion, and subsequent loss of economic value.

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FORAMINIFERA AS ENVIRONMENTAL INDICATORS To obtain a better appreciation of Foraminiferida, and their significance in oceanographical and paleontological studies, we shall touch upon topics that may seem to be out of place in a section dealing with coral reefs. However, since foraminifera have played the key role in the unraveling of paleoclimates and paleoceanographic conditions, it seems appropriate, even though a deviation from standard practice, to discuss the various techniques used to extract environmental information from deep-sea CaC0 3 deposits, especially foraminiferal tests. Studies on Foraminiferida have increased dramatically since the early 1970s with the realization of their value as tools in exploration geology (Hinte 1978) and as indicators of paleoclimatic (Blow 1969; Molfino et al. 1982; Barmawidjaja et al. 1989a, b; van der Kaars 1991) and paleoceanographic conditions (Vincent and Berger 1981; Ganssen et al. 1989; Charles and Fairbanks 1992; Whitman and Berger 1992). However, up to now there has been very little research conducted on the biology or ecology of benthic and pelagic foraminifera in Indonesia (Adisaputra 1989; Kleijne et al. 1989; Troelstra and Kroon 1989). Species Associations The co-occurrence of particular species groups of planktonic and benthic foraminifera as well as the presence of other planktonic microfossil groups (e.g., calcareous nanoplankton and radiolaria) is often used as an effective tracer of various water masses (Adisaputra 1989; Troelstra and Kroon 1989), as well as of past and recent environmental conditions (Troelstra et al. 1989). Close association of various species groups of foraminifera with specific water masses has proved in the past to be a very useful tool for paleoceanographic reconstructions (Be and Tolderlund 1971; Hemleben and Spindler 1983). This is possible because many species have specific environmental and ecological requirements. Temperature and nutrient concentrations are the two dominant environmental factors that seem to regulate the presence or absence of certain pelagic species groups. With regards to the distribution of benthic foraminifera, the primary environmental factor involved is depth (van Marie 1989). Troelstra and Kroon (1989) have demonstrated that during the Northwest Monsoon (i.e., period of downwelling), planktonic foraminifera (28 known species) in eastern Indonesia can be grouped into five distinct (statistically derived) "associations", each characteristic of a specific region (fig. 9.11). Note that, unlike other studies which use bottom sediment samples to identify distribution patterns of pelagic foraminifera, Troelstra and Kroon (1989) used surface water samples. The species associations identified by Troelstra and Kroon (1989) seem to be reflecting subtle changes in environmental parameters (e.g., sea surface temperatures and nutrients) that occur between the western Flores Sea and the eastern regions of the Banda Sea and Timor Trough. Adisaputra (1989), using the sediment-analysis approach, demonstrated that planktonic foraminiferal communities in the Flores Sea differ considerably from those found in the Lombok and Savu Basins (table 9.6). While some of these differences can be attributed to the depth of the lysocline and the carbonate compensation depth (see following discussion), Adisaputra (1989) concluded that the distribution of Neogloboquadrina dutertrei, which is the dominant planktonic foraminiferan in the Flores Sea, is

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Figure 9.11. Recent planktonic foraminifera of eastern Indonesia from sediment cores taken in Flores, Savu and Lombok Basins. A) Globorotalia menardii; B) Globorotalia tumida; C) Pulleniatina obliquiloculata; D) Neogloboquadrina dutertrei; E) Globigerinoides cyclostomus; F) Globigerinoides ruber, G) Globigerina bulloides. Photo courtesy of M. K. Adisaputra,

Pusat Pengembangan

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largely a function of hydrological conditions. Using faunal analysis (i.e., changes in species assemblages) on sediments from a deep-sea core taken in the Molucca Sea, Barmawidjaja et al. (1993) were able to demonstrate that the continual presence of Neogloboquadrina dutertrei (herbivorous foram) since the last glacial suggests that the Deep Chlorophyll Maximum (DCM) layer has been a permanent feature in the region for the past 27,000 years. Of the 16 planktonic species recorded in the study, four species, namely Globorotalia menardii, Pulkniatina obliquiloculata, Neogloboquadrina dutertrei, and Globigerinoides ruber, were found in all stations above the lysocline. The planktonic foraminifera fauna in table 9.6 is characteristic of oceanic regions with well-developed DCM layers (Ravelo et al. 1990), and is very similar to the faunal composition of the north Molucca Sea deep-sediment core (Barmawidjaja et al. 1993). In a parallel study, Troelstra and Kroon (1989) found that the Globigerinoides ruber, Globigerinoides sacculifer and Globorotalia menardii association (group composition 29%-58%, 12%-29%, and 7%-24%, respectively) seems to be characteristic of the open central Banda Sea surface waters, while the Globigerinoides ruber, Globigerinoides sacculifer and Neogloboquadrina dutertrei association (group composition 40%60%, 8%-22% and 5%-16%, respectively) is characteristic of waters along the southern tip of Sulawesi and northeast Timor. The Banda Sea association, which is dominated by G. sacculifer and G. ruber, is characteristic of oligotrophic (i.e., lownutrient) conditions known to exist in the region during the Northwest Monsoon, when surface downwelling predominates as a result of strong northwesterly winds (Wyrtki 1961). It is expected that surface distribution patterns of planktonic foraminifera will change significantly during the Southeast Monsoon, when upwelling generated by the strong southeasterlies predominates in the region. Troelstra et al. (1989) have in fact shown this to be true, by comparing surface abundance of planktonic foraminifera with the abundance of their tests from the surface layers of deep-sea sediment samples. The high abundance of Globigerina bulloides in the surficial layers of deep-sea sediments taken in the Seram and Tanimbar Troughs, two well-known regions of upwelling, and its virtual absence from the surface zooplankton community during the non-upwelling period (i.e., Northwest Monsoon), demonstrates conclusively that G. bulloidesis a good indicator of paleoupwelling conditions (Troelstra et al. 1989). The high abundance of G. bulloides tests in deep-sea sediments indicates that productivity during the upwelling season must be considerable. Subpolar Forams in Indonesia Surprisingly, the subpolar, dextrally coiled, planktonic foraminifera Neogloboquadrina pachydermia has recently been recorded in the tropical waters of the Indonesian Archipelago (Barmawidjaja et al. 1993). This interesting discovery was made during the Snellius-II Expedition to the northern Molucca Sea. The main objective of the study was to obtain climatic proxy data for the reconstruction of climatic conditions in the northern Molucca Sea region during the last glacial period. Neogloboquadrina pachyderma was recovered from a deep-sea sediment core taken in the northeast sector of the Molucca Basin, about 50 km northwest of Halmahera. The sediment core was taken from a depth of 3510 m, and consisted of 533 cm of greenish clays with a rich and well-preserved foraminifera fauna as well as radiolarians and diatoms (Barmawidjaja et al. 1993). One of the most exciting discoveries was the

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Figure 9.12. Map of Flores and Savu Seas showing locations reported in table 9.6.

Table 9.6. Abundance of some planktonic foraminifera tests in deep-sea sediment samples in Flores Sea: A (1537 m), B (1911 m), C (1500 m); Savu Sea: D (3295 m), E (3305 m), F (2514 m) and Lombok Basin (Indian Ocean) G (2085 m), H (4132 m) as for figure 9.12. Numbers in parenthesis are sample depths. Values in table represent percentages of the most common planktonic foraminifera in relation to total sediment composition. Data compiled from stations lying above the lysocline. Species Globorotalia menardii Globorotalia tumida Globorotalia ungulata Pullenlatina obliquiloculata Neogloboquadrina dutertrei Globigerinoides ruber Globigerinoides sacculifer Globigerinoides immaturus Globigerina bulloides Globigerina calida Sphaeroidinella dehiscens Orbulina universa Globigerinella siphonifera Total

Flores Sea Savu Sea A B C D E 3.3 1.7 20.3 2.5 7.4 0.1 3.1 0.2 0.5 0.3 0.1 1.2 0.2 0.1 4.4 1.2 10.1 1.2 0.7 9.3 2.2 21.1 0.8 1.5 3.0 0.9 3.1 0.3 1.3 0.3 0.1 1.8 0.5 0.3 0.3 1.8 0.2 0.1 1.3 1.6 0.4 0.1 0.2 0.1 1.8 0.7 8.6 0.4 24.0 8.9 71.1 6.4 12.3

F 18.6 1.6 2.0 2.8 1.5 0.9 0.1 1.1 0.6 0.8 0.6 30.6

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0.1

0.2 40.0

0.3 4.8

Note on the modification of Adisaputra (1989): Globigerinoides sacculifer and G. sacculifer are the same species. Globigerinella (= Hastigerina). Globigerinella siphonifera and G. aequilateralis are synonyms, with G. siphonifera taking precedent. Modified from Adisaputra

1989.

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great abundance of dextrally coiled N. pachyderma in sediment layers dated to the last glacial (i.e., 28-14 ka). The presence of N. pachyderma in the sediment core is rather unexpected, since it is a cold-water species. It was probably brought in by the intermediate depth current. Thus, its presence has some important paleoclimatic and paleoceanographic implications. According to Barmawidjaja et al. (1993), its occurrence in the northern Molucca Sea is most likely related to the nutrient-rich DCM layer, since it feeds exclusively on phytoplankton (i.e., a herbivore), as does its warm-water morphotype Neogloboquadrina dutertrei. N. dutertrei is a dominant planktonic foraminifera in the eastern Indonesian waters, as is indicated by the sedimentary record (Adisaputra 1989). D e e p Chlorophyll Maximum Layer The Deep Chlorophyll Maximum layer is an ubiquitous feature of many tropical regions throughout the year, and is characterized by a significant increase in chlorophyll concentrations (i.e., up to a factor of 30). The high abundance of Neogloboquadrina dutertrei since the last glacial has been suggested as a proxy record for the presence of a well-established DCM layer in the eastern regions of the archipelago (Barmawidjaja et al. 1993). The depth of the present-day DCM layer in the northern Molucca Sea has not been established, but it is most likely associated with the nitracline (i.e., zone of maximum change in nitrate concentrations), and near the primary nitrite maximum (Parsons et al. 1984). In general, the DCM is located near the bottom of the euphotic zone, and, therefore, it should be located at depths of less than 100 m. Based on our reconstruction of vertical depth profiles (Oceanographical Cruise Reports No. 23 and 36) for temperature and nitrate during both upwelling and non-upwelling periods in the northeast Molucca Sea, it is suggested that the DCM layer is most likely located between the depths of 150 m to 75 m (fig. 9.13). Chlorophyll sampling of sufficient refinement to properly identify the DCM layer, which is only a few metres thick, is not currently available from this region. However, Gieskes et al. (1989) were able to demonstrate that a well-defined DCM layer does indeed develop during the Northwest Monsoon in the eastern regions of the Banda Sea. The DCM layer becomes established in the thermocline during the non-upwelling period at a depth of 60 to 80 m (Gieskes et al. 1988). Figure 9.14 illustrates the development of a DCM layer in the east Banda Sea at Station 9 (04°39.2' S and 130°30.3' E) during February 1984 (Northwest Monsoon) and August 1985 (Southeast Monsoon) (Wetsteyn et al. 1990). Note, that the DCM layer has been observed to "spill" over offshore reefs as a result of strong tidally-induced currents or internal waves. The frequency of these events and their ecological significance has not been investigated. The apparent upwelling event, indicated by the shallowing of the isopycnals (fig. 9.14A), is most likely associated with the intensification of the Mindanao Current (gyre) during the Southeast Monsoon. Strong southerly (i.e., from the south) monsoon winds that predominate in this region from July to October create strong northward-flowing surface currents from the northeast Molucca Sea with a net outflow into the Pacific Ocean (Wyrtki 1961). The net surface outflow from the northeast sector of the Molucca Sea is then balanced by upwelling of deeper nutrient-rich water along the west coast of Halmahera.

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Figure 9.13. Temperature and nitrate vertical depth profiles for Station No. 11, Moluccas (Moluccas) Sea. Data compiled from Oceanographical Cruise Reports No. 23 and 36. Note the upward displacement of both the thermocline (A) and nitracline (B) during the upwelling period (Southeast Monsoon) in August 1977. The position of the Deep Chlorophyll Maximum (DCM) layer is suggested by the shaded areas. Source: Institute of Marine Research 1977; National Institute of Oceanology 1982.

The analysis of the Molucca Sea sediment core revealed that at about 14 ka (at 270 cm depth), there was a pronounced shift in community composition of the planktonic foraminifera. The abundance of Neogloboquadrina pachyderma and Neogloboquadrina dutertrei (sinistral) increased, while there was a marked decrease in the abundance of the oligotrophic species group Globigerinella aequilateralis, Globigerinoides ruber and Globigerinoides sacculifer. Barmawidjaja et al. (1993) suggested that the marked increase of the two phytoplankton-feeding species with a concurrent reduction in the oligotrophic species group, indicates that the glacial DCM layer in the north Molucca Sea may have been considerably more productive than at present. However, the presence of the dextrally coiled Neogloboquadrina pachyderma is more problematic, and its origins remain highly speculative. New research efforts are required to solve some of these fascinating questions. Morphological Indicators The calcareous tests of planktonic and benthic foraminifera as well as a host of other calcifying marine organisms are excellent recorders of environmental con-

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Figure 9.14. Deep Chlorophyll Maximum layer in the eastern Banda Sea (Station 9, 04°39.2' S and 130°30.3' E, Snellius-ll Expedition) during: A) February 1985; and B) August 1984. Note that the DCM is associated with the primary nitrite maximum. Source: Wetsteyn et al. 1990.

ditions in which the animal or plant lived, and are, therefore, of particular interest to the paleontologists. One of the most pressing environmental concerns of the industrialized and non-industrialized world is the anthropogenically-induced global climate change, which according to many scientists is now inevitable (Houghton et al. 1990). The knowledge of past climates, paleoceanographic conditions and the reconstruction of ancient atmospheric C 0 2 concentrations are essential to our understanding of the environmental changes that may occur, and foraminifera have been in the forefront of this pursuit. These organisms have been used extensively in the retrieval of climatic data from the geologic record, which is essential for validation of current models in which oceanic temperatures are of paramount importance (Becketal. 1992).

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Test Porosity Foraminifera have proved to be excellent indicators of environmental conditions, mainly because of their specific environmental and ecological requirements (Kennett 1976), and their relative ease of handling. Many groups of foraminifera exhibit well-defined intraspecific phenotypic differences (i.e., variations in test porosity, texture, thickness) which allow for separation of distinct sub-populations, each with subtly differing environmental requirements (Kennett 1976). One simple feature of foraminiferan test morphology that does not require elaborate analysis (only a simple microscope), and which has been of great assistance to paleontologists, is test porosity. Test porosity simply refers to the number of pores in the test; however, the value is presented as mean porosity which is the percentage of open pore area per unit area of test (Murray 1991b). This measure takes into consideration not only the number of pores, but also the pore size. Interspecific differences in the porosity of planktonic foraminifera are quite pronounced (Murray 1991b). These differences become especially apparent when viewed from a latitudinal perspective (i.e., temperature) . Thus, we find that test porosity is low in higher latitudes and high in low latitudes, which roughly correlates with temperature differences. For example, Neogloboquadrina pachyderma, which is a high-latitude form (i.e., subpolar), has a mean test porosity of about 2%, while Globigerinoides sacculifer, which is abundant in the Indonesian waters (i.e., tropical), has a mean porosity of 18% (Murray 1991b). However, there are also significant intraspecific differences in test porosity which are related mainly to ontogeny (i.e., history of development). Unfortunately, only a few life history studies on planktonic or benthic foraminifera have been conducted thus far (Brummer et al. 1987), and none have yet been attempted in Indonesia. Intraspecific differences in test porosity have also been used to differentiate between cold-water and warm-water pelagic species (Be 1968,1977; Be et al. 1973). In their study of the Indian Ocean foram Orbulina universa, Be et al. (1973) demonstrated that populations from low latitudes (i.e., the equatorial Indian Ocean) had a significandy higher porosity than foraminifera in higher latitudes. O. universa has been reported from the Savu Sea (Adisaputra 1989), Segara Anakan, South Java (Subardi et al. 1989) as well as from Ambon Bay (Hermanto and Suhartati 1989), however, no information on its test morphology is available. An explanation for intraspecific differences in test porosity has so far been very elusive; thus while there are demonstrable effects from such factors as temperature, light, salinity, nutrient supply and seawater density, the development of test porosity and its functional role are still a mystery (Be 1977; Hemleben et al. 1985; Hemleben etal. 1987). Test Coiling A number of foraminiferal families (planktonic and benthic) exhibit pronounced test coiling. A trochospiral condition exists when foram test chambers are spirally coiled, evolute (i.e., unfolded and where all whorls are visible) on one side and involute (i.e., tightly coiled with only the final whorl visible) on the opposite side (e.g., Trochamminidae, Globorotaliidae, Eponididae, Nonionidae). The coiling of the test can be either dextral (i.e., clockwise, coiling from left to right) or sinistral (i.e., anticlockwise, coiling from right to left). Brummer and Kroon (1988) have demonstrated experimentally that coiling in foraminiferal tests is to a large extent genetically determined. However, this does not explain why a particular species

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would have both dextral and sinistral growth-forms; the functional significance remains elusive. Of particular interest, however, is the fact that some species have distinct geographic populations whose members are predominantly trochospiral in one direction only, while a population in another geographic area exhibits an opposite coiling direction. The cold-water planktonic foraminifera Neogloboquadrina pachyderma has become a classical example of this phenomenon (Arikawa 1983). N. pachyderma is distributed widely throughout the polar and subpolar regions. The polar and subpolar populations are at present most likely reproductively isolated from one another. In cold polar regions, coiling in N. pachyderma is generally sinistral, while in the slightly warmer subpolar regions the coiling is dextral. One is tempted to evoke temperature as the obvious controlling factor, and attempt to use this morphological feature in paleoclimatic analyses. However, Murray (1991b) suggested a cautious approach to the application of coiling ratios in paleoecological or biostratigraphic studies.

T H E U S E OF S T A B L E ISOTOPES With advances in modern technology, new methods were developed that allow us to extract a wealth of stored environmental information from the foraminiferan tests, and other CaC0 3 -secreting organisms. The two techniques that have become indispensable to paleontologists, in their reconstructions of paleoclimates, are the analyses of stable (i.e., not radioactive) oxygen (8 O) and carbon (8 C) isotopes in calcareous skeletal material. There has been a recent surge in the use of other elements, such as nitrogen (8 15 N), sulfur (834S), strontium and calcium ratios (Sr/Ca), boron (8 B) and hydrogen ( H and D - deuterium). The use of stable isotope analysis to solve various biogeochemical problems has increased dramatically in the past decade, mainly because stable isotope data are now used by both physical scientists (e.g., oceanographers, geologists) (Peterson and Fry 1987; Salomons and Mook 1987; Wu and Berger 1991; Beck et al. 1992; Lehman and Keigwin 1992; Farquhar et al. 1993; Plank and Langmuir 1993; Spivack et al. 1993) as well as by biologists in purely process-oriented biological or ecological studies (e.g., ontogeny, food-web analysis) (Eichler 1966; Weber 1974; Land and Lang 1975; Black and Bender 1976; Goreau 1977; Estep and Dabrowski 1980; Schroeder 1983; Peterson et al. 1985; Ehleringer et al. 1986; Rounick and Winterbourn 1986; Muscatine et al. 1989; Monteiro et al. 1991; Durako and Hall 1992; Murphy and Kremer 1992; Cerling et al. 1993). Stable isotope techniques are possible because each element has at least two stable isotopes (each with slightly different physicochemical properties), one of which is always present in greater abundance than the other. As the elements cycle through the biosphere, they undergo numerous chemical reactions that alter the isotopic composition (i.e., fractionation) of the elements in a predictable way (Peterson and Fry 1987), and these differences can now be measured with great accuracy. Stable Oxygen Isotopes Oxygen (O2) concentrations in seawater have long been used by physical oceanographers to trace the origins of various oceanic water masses, and together with

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salinity to determine their mixing rates (Broecker 1974). Oxygen became a powerful oceanographic tool when it was found that the two stable isotopes (i.e., heavy 18 O and light 16 O) that comprise the oxygen element have subtly-differing physicochemical properties. Water molecules bearing 18O (i.e., heavier isotope) evaporate much slower from the sea, but condense faster than water molecules containing 16 O; thus a "fractionation" of oxygen occurs as the element cycles through the atmosphere. This has a great influence on the isotopic composition of rainwater as well as the water masses from which the water initially evaporates. Oxygen-18 is present in seawater in much smaller quantities than 1 6 0; in fact, 99.8% of all water molecules in the oceans bear 16O and only 0.2% bear 18O. Various water masses in the world's oceans are characterized by specific 18O/ 16O ratios, and a generalization can be made that surface waters in high latitudes contain lower 18O content than those in the equatorial regions (Broecker 1974). In other words, warm tropical water, because of its excess evaporation, becomes enriched in 18O, while high-latitude cold water becomes enriched in 16O. Another important feature is that the 18O/ 16O ratio is also salinity-dependent. High-latitude water masses characterized by low 18O content also have low salinities (Whitman and Berger 1992), while high-salinity equatorial water masses are enriched with 18O. The differences in the physico-chemical properties of the two oxygen isotopes have provided a powerful tool for extracting paleoenvironmental information from the fossil record ever since the technique was first introduced (Emiliani 1955). Oxygen isotope analysis of the calcareous remains of foraminifera, corals and other calcifying organisms (e.g., Halimeda) has been applied widely in environmental and ecological studies (Patzold 1984; Margosian et al. 1987; Siegrist and Randall 1989; Klein et al. 1992; Oba et al. 1992). All calcifying aquatic organisms must incorporate stable isotopes into their CaCO3 skeletons during the calcification process, the source of the oxygen being the C 0 2 as either HC03~ or C03~. Among the various marine calcifying organisms, the isotopic ratio in foraminiferal and coral CaC0 3 is usually close to isotopic equilibrium with seawater (Urey 1947; Emiliani 1955, 1978; Wefer and Berger 1991). Note that in foraminifera and corals, some degree of 8 18O fractionation occurs. One explanation for generally low 8 18O values (relative to equilibrium) in scleractinian skeletons is that the HC0 3 " used during the calcification may be in isotopic disequilibrium with seawater (McConnaughey 1989a,b), Spero and Lea (1993), working with Globigerinoides sacculifer under laboratory conditions, showed that shell 8 1 8 0 values vary as a function of symbiont photosynthesis, but do not vary with ontogeny. Nontheless, the 8 18O technique is very useful, since any changes in the ambient (i.e., seawater) isotopic composition will be reflected in the isotopic composition of the CaC0 3 deposited at the time. Thus, the skeletons (e.g., foraminiferan tests, coral coralla, bivalve shells, cephalopods shells, etc.) in fact become recorders of the environmental (e.g., seasonal temperatures) as well as physiological changes that occur during the lifetime of the organism (Romanek et al. 1987). Based on samples recovered from deep-sea sediment traps set in the Sargasso Sea, Deuser et al. (1981) were able to demonstrate that in Globigerinoides ruber, which is the most frequently used species for isotopic determinations in Indonesian studies (Ganssen et al. 1989; Gayet et al. 1990; Barmawidjaja et al. 1993), deposition of CaC0 3 occurred close to isotopic equilibrium with surface seawater. Experimental laboratory studies with Globigerinoides sacculifer and Orbulina universa demonstrated that CaC0 3 deposition indeed occurs under

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isotopic equilibrium with seawater, and that changes in oxygen isotopic values are directly related to temperature, provided all other variables remain constant (Erez and Luz 1983; Bouvier-Soumagnac and Duplessy 1985). Thus these species are useful biomarkers. There are, however, exceptions to the equilibrium model, as, for example, in Globigerina bulloides, an indicator of upwelling conditions in the Banda Sea, whose tests in the Sargasso Sea (North Atlantic), are known to be deposited in isotopic disequilibrium with surface waters (Deuser et al. 1985). Thus analysis of oxygen isotopes in this species would be of limited value without additional data. The 1 8 0 / 1 6 0 isotopic ratio in the CaCO3 skeletons of modern foraminifera is primarily dependent on the ambient seawater'temperatures (Urey 1947), salinity (Craig 1965; Tan and Strain 1980; Whitman and Berger 1992) and the isotopic composition (i.e., 18O/ 16O ratio) of seawater from which the carbonate is precipitated (Patzold 1984; Romanek 1987; Whitman and Berger 1992). The analysis of 18 16 O/ Oin CaCO 3 skeletal material is somewhat complicated, because the temperature effect may be masked either by oxygen isotopic changes due to variations in physiological activity, especially in organisms associated with photoendosymbionts (e.g., foraminifera, corals, giant clams), or by salinity differences. However, under constant salinity and 1 8 O/ 1 6 O ratio of seawater (i.e., most oceanic coral reef environments), temperature has an overriding influence on the incorporation of oxygen isotopes into the foraminiferal tests, and other CaC0 3 skeletons. In general, as seawater temperatures increase, the 18O/ 16O ratio in the carbonate decreases correspondingly and vice versa. This is very useful information, since using the isotopic ratios, from foraminiferan tests, for example, we can calculate the seawater temperatures under which the CaCO3 was deposited. Thus, the tests of Globigerinoides ruber, which has a wide geographic distribution, will contain more 18O if the formation of the test occurred in a cold-water environment (e.g., upwelling area, or winter months), while the tests will have lower 18O content if the tests were formed in a warm-water environment. Indeed, this relationship was observed in plankton tow studies from Bermuda (Williams et al. 1981). However, great caution is needed to interpret oxygen isotopic data in organisms such as planktonic foraminifera, since throughout their different life-cycle stages they require different environmental conditions (i.e., depth-temperature related), and thus the continuous deposition of CaC0 3 could proceed under very different temperature regimes, which is reflected in the isotopic signature. For example, most juvenile planktonic foraminifera (especially symbiotic species) mature in the warmer surface waters (0-30 m), and upon reaching sexual maturation they migrate into deeper-cooler waters just prior to gametogenesis (Murray 1991b). It is now believed that they migrate down to the top of the seasonal thermocline and DCM (H. Spero, pers. comm.). As expected, isotopic ratios in the tests of these species will reflect the extent of the vertical migration, since it is registered in the CaC0 3 by 18O enrichment. Obviously, knowledge of their ontogeny is essential to unravel these life-cycle-related complexities in the isotope record (Murray 1991b). Ice Volume Effect Another factor that needs to be mentioned when discussing the 8 1 8 0 technique is the magnitude of the ice effect (Emiliani 1955, 1966). This effect is quite considerable, since large volumes of water are moved from ice to sea during the glacial/interglacial cycle, causing large-scale sea level fluctuations (Milliman and

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Emery 1968). The maximum sea-level drop during the last glacial period in regions outside the archipelago have been estimated between 120-130 m below present-day sea level (Milliman and Emery 1968; Chappel and Shackelton 1986; Fairbanks 1989). Recent studies on sea-level fluctuations in the Indonesian Archipelago indicate that during the last glacial maximum, the lowest sea level stand was somewhere between 148(±16) m and 106(±11) m below present-day sea level (Hantoro 1992). This implies that during the last glacial maximum, about 3.5% of ocean volume was stored in polar ice caps and continental glaciers. According to Whitman and Berger (1992), 8 18O values of polar ice were between -38%o and -34%o (i.e., ice is greatly depleted in 1 8 O). During glaciation, 18O preferentially stays in oceans (it is fractionated), thus increasing the seawater 1 8 0 / 1 6 0 ratio (8 1 8 0 increases), and as expected, the foraminifera that happened to live during this period will have tests that are correspondingly enriched in the heavier isotope (i.e., 8 18O increase). During deglaciation, when the glaciers release large volumes of water, the 8 18O values decrease correspondingly. The difference in 8 18O between the glacial and interglacial periods (i.e., ice effect) in deep-sea sediments is 1.2%o, which means that for every 10 m change in sea level the 8 1 8 0 values change by 0.1 %o (Whitman and Berger 1992). Thus the 8 18O sediment record has also been used as an important sea-level marker. Foraminifera have stored all this information in their CaCO3 test. Measurement All stable isotopic ratios are measured with a mass spectrometer (Schneider and Jones 1992), and the results are by convention presented in delta (8) values, which are expressed as deviations per mill (i.e., %o) relative to an international standard PDB (Peedee Belemnite CaCO 3 ). The delta values are calculated as: 5

i80%o

=

|" (S 18 Q / 8 1 6 Q ) sample - (S 18 Q / 8 1 6 Q ) s t a n d a rd ~| 18

x

1QQ0

16

L (S 0 / 8 0 ) standard J The 8 values are thus ratios of heavy and light isotopes in the sample, compared to a known standard. The standard used depends on the analysis being done. For example, in 8 18O analysis of foraminiferan tests or coral skeletons (i.e., calcite or aragonite) the standard is PDB. It is useful to remember that any increase in ' 8 ' value implies a corresponding increase in the heavier isotope. The 8 18O is now a widely used technique in paleontology, since 8 18O isotope analyses of deep-sea cores containing calcareous organisms can reveal information on seawater temperatures at the time when the animal was alive and depositing the skeleton, which may have been millions of years ago. One equation commonly used to convert the 8 18O values to temperatures is the paleotemperature equation developed by Epstein and Mayeda (1953): T(°G) = 16.5 - 4.3(8 18 O c - 818Ow) + 0.14(8 18 O c - 818Ow)2 where 8 18Oc and 8 18Ow refer to shell and water samples, respectively. However, the isotopic composition of seawater varies with salinity, which is a function of evaporation, precipitation, freezing, mixing, and fresh-water runoff. To circumvent some of these problems, reconstruction of paleotemperatures has relied heavily on the use of foraminiferal shells from deep-sea sediments. One limitation of the present

THE USE OF STABLE ISOTOPES

409

8 1 8 0 technique is that the temperature interpretation of foraminiferal isotopic data extracted from deep-sea sediments is based largely on our knowledge of modern foram physiology, and their response to temperature changes under laboratory conditions. However, new techniques are being developed that may circumvent some of these problems. One of the most promising is the strontium/calcium ratio technique (Beck et al. 1992), since the incorporation of the two elements into the CaCO3 skeletons (i.e., aragonite) is a function of temperature, but only a weak function of salinity (Beck et al. 1992). Using this technique, Beck et al. (1992) suggested that about 10,200 years ago the average sea surface temperatures in the southwestern Pacific Ocean were about 5°C cooler than they are today, and that the tropical belt was much narrower. Sr/Ca studies using massive skeletons of Pontes spp. have apparently been conducted in Indonesia; however, the results are not available. Use of 8 1 8 0 in Ontogenic Studies To digress a bit, we should point out that oxygen isotopes can be very useful in tracing life history (i.e., ontogenic) changes in habitat depths of other organisms which may be difficult to observe in their natural habitats. A prime example is Nautilus pompilius (Cephalopoda), a deep-water, reef-associated cephalopod which is frequentiy sighted in deeper regions (>50 m) of the reef slope on many reefs in the Pacific region (fig. 9.15). The population demography of Nautilus in Indonesia is still a mystery, yet its shells are sold in great quantities in almost all souvenir and handicraft stores in the country, even though it is a fully protected species. The fossil record of the extant Nautilus dates to the Oligocene; however, the Subclass Nautiloidae (i.e., cephalopods with calcareous external multichambered siphunculate shells) first appeared in the late Cambrian, and their basic body plan has not been altered since the Ordovician (O'Dor et al. 1993). Nautilus jav'anus from the late Miocene is the only fossil cephalopod described from Java (van Bemmelen 1949). There are a number of extant species, the most common being Nautilus pompilius, which is a true living fossil (Saunders 1987). This deep-water invertebrate is known on occasions to visit the deeper parts of the lower reef slopes, but the adults usually stay at depths below 100 m, and none have so far been captured at depths of less than 50 m (Obaetal. 1992). The use of 8 18O was instrumental in unraveling the mysteries of its ontogenic changes in habitat depths. We should also point out that successful hatching of aquarium-held Nautilus belauensis has greatly expanded our understanding of its ontogeny (Carlson et al. 1993). It appears that Nautilus has an exceptionally long incubation period lasting about 12 months. According to Oba et al. (1992), incubation takes place in the shallowest depths of its overall depth distribution, which may, however, vary significandy between different regions depending on other environmental conditions. Oba et al. (1992) were able to demonstrate that nautili in the Philippines (i.e., Taiion Strait) hatch at depths between 100-80 m (24°-28°C). Based on the 8 18O values in nautili shells, it was demonstrated that once hatched, the juvenile N pompilius live at the hatching depth until the deposition of the twelfth septum. The juveniles then move into deeper water where temperatures are about 17°C, and can be as low as 13°C, which is reflected in the sharp increase in

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Figure 9.15. A true living fossil, Nautilus pompilius is a deep-water reef-associated cephalopod that occurs throughout the Pacific from Fiji to Indonesia. Photo courtesy of R. Steene.

8 18O values, from about -0.3%o (or lower) to values roughly between 0.6%o to 1.2%o (Oba et al. 1992). The use of 8 1 8 0 values was also instrumental in determining that Nautilus in Fiji inhabit considerably deeper waters, which was reflected in enrichment in 8 18O in both juvenile (range 0.1 %o to 0.5%o) and adult (range 1.4%o to 2.0%o) nautili. Thus hatchlings and juveniles are found at depths between about 300 to 350 m, while adults live at depths between 450 to 550 m (Oba et al. 1992). The usefulness of the 8 18O technique was further demonstrated by analyzing the isotopic composition of the last septum of mature adults. Micro-sampling taken from the last septum showed regular fluctuations in 8 1 8 0 values of about 0.3%o to 0.5%o, roughly corresponding to differences in 1°-2°C. These regular 8 1 8 0 fluctuations correspond to daily vertical migrations of the species. This has been corroborated with telemetric studies at the University of Papua New Guinea, Motupore Island Research Station (O'Dof et al. 1993).. Vertical movements of Nautilus along the reef face are mostly crepuscular, ascending at dusk to shallower reefal waters in search of food and descending into the deeper, cooler waters at dawn, presumably to conserve energy (O'Dor et al. 1993). Recent isotopic studies, as well as successful hatchings of nautili at the Waikiki Aquarium (Carlson et al. 1993), have provided important clues to the ontogeny of this living fossil. The requirement for warm water for incubation seems to have answered one of the major questions of Nautilus distribution in the tropics (Carlson et al. 1993). Thus, while the species is obviously temperate in character, since it prefers cooler deeper waters, they need the tropical warm water for incubation. Thus, their reproductive strategy has restricted their distribution to the tropical oceans. It is believed that Nautilus has only a limited ability to swim between islands

STABLE C A R B O N I S O T O P E S

411

(Saunders 1987), and thus the absence of a planktonic larval stage during its ontogeny greatly reduces its dispersal ability, and partly explains its rather narrow distributional range.

STABLE C A R B O N ISOTOPES Perhaps of greater value in recent paleontological investigations has been the use of stable carbon isotopes (i.e., 12C and 13C). Carbon-12 is much more abundant than 13 C, representing 98.89% of the carbon pool, while 13C accounts for the additional 1.11%. Studies of 12 C/ 13 C (i.e., 813C derived as for 8 18 0) ratios in animal skeletons (especially foraminifera) have provided valuable information on past carbon cycles, and most importantiy on the productivity of the paleoceans. In 8 G determinations the international standard is Belemnitella americana from the Cretaceous PeeDee Formation (PDB). Unlike the oxygen isotopic ratios, the 8 C is not temperaturedependent. This means that the 8 C value in shells of various calcifying organisms, or coral skeletons, is very close to the ambient values in which the CaCO3 was precipitated, and is thus useful in assessing productivity (i.e., photosynthesis) of paleoceans at the time of CaC0 3 deposition. Indeed, high 8 C values in foraminifera tests have been correlated with high primary productivity of the water masses in which the foraminifera deposited their tests. The 8 C enrichment occurs because during phytoplankton blooms, or in regions with intense primary productivity, phytoplankton preferentially take up 12C, which leads to a substantial enrichment of C in the XC0 2 of surface waters (Ganssen et al. 1989). During photosynthesis 12 C atoms are fixed 1.02 times more rapidly than 13C atoms. Since 813C in CaCO 3 skeletons reflect ambient 813C levels, foraminifera from regions of high primary productivity will have tests usually enriched in 8 C (Ganssen et al. 1989). Carbon isotopes have also been useful in ontological studies of planktonic symbiotic foraminifera. Tests of juvenile symbiotic planktonic foraminifera have been shown to be depleted in 13C (i.e., lower 813C) when compared to larger individuals. The isotopic difference is a result of vital effect related to the photosynthetic activity of the photoendosymbionts (mostly dinoflagellates and chrysophytes). Juvenile foraminifera are usually found in the upper layers of the euphotic zone until maturity, and, as a result, photosynthetic activity of the endosymbionts is high (Spero and Parker 1985), thus producing ample quantities of isotopically light (i.e., C depleted) carbon pool which is used by the juvenile foram in CaCO s deposition. As a result, the tests of all small (i.e., juvenile) symbiotic foraminifera are S13C depleted. Differences in 8 C values and morphological variations in test architecture are powerful tools to distinguish different foraminiferan sub-populations. Perhaps one of the best examples is that of Globorotalia truncatulinoides, whose subantarctic and subtropical populations in the southern Indian Ocean differ in their 8 C signatures and test morphology (Healy-Williams et al. 1985).

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P A L E O C E A N O G R A P H Y AND P A L E O C L I M A T E S Stable isotopes have become important portholes into the distant past, and have provided new insight into the paleoclimatic and paleoceanographic conditions during the last glacial period in the Indonesian Archipelago. It is well-established that, during the last glacial maximum, sea surface temperatures in most equatorial oceanic regions dropped by about 2°-3°C, with a corresponding equatorward displacement of polar fronts (Seibold and Berger 1982). Surface waters of many subtropical and tropical coastal environments (e.g., eastern Indonesia) may have been additionally cooled by the intensification of coastal upwelling. As expected, 8 18O values of planktonic foraminifera (recovered by sediment cores) in many regions are enriched in 1 8 0, reflecting the general cooling of the oceans. Note that the temperature enrichment is in addition to the ice volume effect that globally accounts for about 1.2%o of the difference between the last glacial maximum and the Holocene high sea level stand (Seibold and Berger 1982). Isotopic analysis of deep-sea foraminiferal sediments has provided important clues concerning paleoclimates and paleoceanography since the last glacial. New techniques using foraminiferal boron isotope (8 1 B) extended the record to the early Miocene (about 21 Ma). Spivack et al. (1993), using a foraminiferal boron isotope as a proxy for surface ocean pH over the past 21 Ma, demonstrated that the surface seawater pH from the early to middle Miocene was between 7.4 to 7.5 pH. This measurement was made possible because the isotopic composition of boron in foraminiferal tests is pH-dependent. Since the pH of surface waters is a sensitive measure of its alkalinity and total inorganic carbon concentrations, Spivack et.al. (1993) concluded that total carbon and alkalinity 21 Ma were 0.6 times higher than present-day values, indicating that the C 0 2 concentrations of the Miocene atmosphere were much higher than at present. By late Miocene (7.5 Ma B.P.) surface seawater pH values reached 8.2 and remained so up to the present. The Miocene was a period of extensive coral reef development with a global climate that was much wetter and warmer. It seems that lower pH values (higher atmospheric C 0 2 concentrations) did not have a detrimental effect on coral reef development at the time. Recent reviews on the effects of global climate change on coral reef ecosystems have, however, suggested that, with an increase in sea surface temperatures and atmospheric C 0 2 concentrations, a decrease in CaCO3 saturation state may be expected (Smith and Buddemeier 1992). Smith and Buddemeier (1992) further suggested that decreased rates of calcification associated with the altered CaCO3 saturation state will reduce the ability of coral reefs to keep up with the expected sealevel rise. It is interesting to point out that low sea-surface pH values (i.e., lower CaCO3 saturation state) from early to middle Miocene were associated with rising sea levels; however, since the end of the middle Miocene, the steadily rising sea-surface pH values (Spivack et al. 1993) were accompanied by a major glacio-eustatic lowering of sea level, reaching one of the all-time lows (c. 220 m) at about 6.6 Ma B.P. (Hutchison 1989). Three recent studies in Indonesia are profiled to illustrate the great value of stable isotope techniques as well as to review some exciting new information on the paleoclimate and paleoceanography of the region. In concert, these studies offer an incredible glimpse into the glacial past, and provide an interesting chronological sequence of events since the last glacial maximum. The following reconstruction of

N O R T H E R N M O L U C C A SEA

413

past events would not have been possible without the environmental record that was extracted from the tests of small foraminifera.

N O R T H E R N MOLUCCA SEA The analysis of the Molucca Sea deep-sea sediment core (discussed earlier) revealed that 8 18 0 values during the last glacial maximum were 0.8%o higher ( 18 0 enriched) than would be expected if ice volume effect was the dominate factor controlling the 8 1 8 0 values in planktonic foraminifera (Barmawidjaja 1993). Based on their palynological, planktonic foraminifera and 8 18O results, Barmawidjaja (1993) concluded that the 18O enrichment during the last glacial maximum in the Molucca Sea was mainly a function of higher salinities, and not temperature. Higher salinities in the surface waters were brought about by increased aridity in the region. They hypothesized that increased aridity during the last glacial resulted in higher evaporation rates, relative to precipitation, which decreased the 8 18O values (i.e., 18 0 enrichment). What caused the increased aridity during the glacial maximum remains equivocal; however, Barmawidjaja et al. (1993) suggested that the causative mechanism is most likely a combination of two climatic factors driven by the asymmetrical distribution of glacial ice cover. While they acknowledge that the InterTropical Convergence Zone (ITZC) did not oscillate as widely (i.e., a reduction in the northern extent) as it does today, it was not the main cause for the increased aridity. To explain this phenomenon, Barmawidjaja et al. (1993) proposed that the most plausible explanation for the increased aridity during the last glacial maximum lies in the possible disruption of the El Nino-Southern Oscillation (ENSO) system. It seems that during deglaciation, the ENSO system resumed its course, and the climate in the eastern regions of the archipelago became progressively wetter (fig. 9.16). The appreciable and rapid reduction in 8 18O values from about 14 ka to 8 ka (fig. 9.16) corresponds to the last deglaciation, and the rapid melting of the continental glaciers and polar ice caps. These results agree well with those of Ganssen et al. (1989), who arrived at similar conclusions. The northern Molucca Sea deep-sea core (Barmawidjaja et al. 1993) provided a wealth of information on the climatic conditions that existed in the region during the last glacial maximum. One of their most significant findings, from the coral reef perspective, is that surface seawater temperatures in the northern Molucca Sea during the last glacial maximum (i.e., 18,000 ka) were at most only 1°C cooler than at present, and that the 1 8 0 enrichment was mainly salinity-driven. This suggests that glacial sea surface temperatures were basically within the optimal range for corals and coral reef development, providing the first conclusive evidence that temperature fluctuations during the Plio-Pleistocene were not a significant factor affecting coral distribution and coral reef development in the archipelago.

414

FORAMINIFERIDA

Figure 9.16. Oxygen (8180) and carbon (513C) isotope record for Globigerinoides ruber from the northwestern Moluccas Sea plotted against time. Original data modified by smoothing procedure to enhance long-term patterns (i.e., glacial - interglacial differences). Time scale is refined chronology based on AMS radiocarbon (14C) dates of monospecific samples of Globorotalia menardii-tumida. Source: Barmawidjaja et al. 1993 and Ganssen et al. 1989.

MAKASSAR S T R A I T Using foraminiferal biomarkers (species assemblages), and 8 1 8 0 of foraminiferan tests from sediment cores taken in the Makassar Strait, Gayet et al. (1990) were able to reconstruct paleoclimatic conditions in the western region of the archipelago during the past 15.5 ka. The results provide additional support to earlier studies demonstrating that the last deglaciation was interrupted by the Younger Drays event (Fairbanks 1989), which is reflected in 8 18O enrichment of planktonic foraminifera (Ganssen et al. 1989; Linsley and Thunell 1990). Following the glacial maximum at about 18.8 ka, the abundance of planktonic foraminifera as well as surface water productivity in the Makassar Strait were below present-day levels. Most of

BANDA SEA

415

the sediments in the Makassar Strait at the time were allochthonous, originating from Kalimantan via the Mahakam River discharge. The 8 18O values of planktonic foraminiferal tests showed that from 15.5 ka to 13.3 ka there was a sharp discontinuity in the 818Ovalues from -0.5 to -1.1 %o, corresponding with rapid meltdown of the polar ice caps. These results have been corroborated by a more recent study in the northern Molucca Sea (see fig. 9.16) (Barmawidjaja et al. 1993). From 13.3 ka to about 10 ka melting of the polar ice caps was slowed, and more stable conditions prevailed. Increased productivity in the surface waters (i.e., higher abundance of planktonic foraminifera) was most likely associated with significant weakening of the water column stratification, which allowed greater mixing between surface and deep water masses. Greater vertical mixing facilitated by the weakening of the water column stratification may have resulted in higher oxygen concentrations of the deep water masses, which increased from about 5.0 mg.l" to about 5.7 mg.l" by 13.3 ka (Gayet et al. 1990). Since about 8.2 ka, circulation in the Makassar Strait seemed to have slowed down, mainly as a result of the reestablishment of a relatively strong stratification during the Holocene. As a result, there has been a measurable reduction of oxygen in the bottom water masses (Gayet et al. 1990).

BANDA SEA During the Snellius-II Expedition, an interesting study was conducted in the Banda Sea region to reconstruct the conditions which existed during the last deglacial period between 15 ka to 8.2 ka. Using 8 18O techniques, Ganssen et al. (1989) were able to show that the last deglaciation did not lead to lower sea surface temperatures in the Banda Sea region, but that it had a considerable influence on the regional weather patterns. As expected, climatic conditions and oceanic circulation patterns 10,000 years ago were much different. Palynology of deep-sea sediment cores revealed that the climate in eastern regions of the archipelago during deglaciation was much wetter, while the upwelling intensity increased significantly, resulting in increased productivity of surface waters, reflected in higher 8 C values (fig. 9.16). Their 818O record of the core clearly demonstrated that deglaciation was interrupted by the Younger Drays event. There appears to be a significant difference in the sensitivity of the marine and terrestrial systems to climatic changes. The response of terrestrial flora (i.e., increase in spore/pollen ratio) to the intensification of the monsoonal system (i.e., wetter climate) began about 1300 years earlier than the corresponding increase in primary productivity (higher 8 C - see fig. 9.16) of surface waters (Ganssen et al. 1989). However, as illustrated in figure 9.16, the increase in primary productivity since about 9 ka has been remarkable. The driving mechanism behind the high productivity of the surface waters in the eastern regions of the archipelago is upwelling, driven by the monsoonal system. What we learned from the tests of planktonic foraminifera is that the disruption of the monsoonal system will ultimately lead to a reduction in marine productivity. More importantly, however, is that weakening of the present monsoonal system will result in rapid climatic changes in the eastern regions of the archipelago. This is clearly evident from the recent ENSO events, which were associated with long periods of drought, conditions that may have existed during the last

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glacial. The anticipated climate change (Houghton et al. 1990) is a serious concern for Indonesia, since the most likely scenario will be a major disruption (i.e., weakening) of the monsoonal system. The paleoclimatic information extracted from the deep-sea sediment cores indicate that, if this occurs, a change to more arid conditions can be expected, perhaps within the life span of the next generation.

CONTRIBUTION TO M A R I N E SEDIMENTS Large areas of the ocean floor are covered by pelagic biogenic sediments, consisting mainly of calcite tests of holoplanktonic foraminiferans, such as the Superfamily Globigerinacea (e.g., Globorotalia, Globigerina, Globigerinoides, Globigerinella). Globigerinoides are a cosmopolitan group which has, however, not been studied in Indonesian waters even though they are a significant component of the zooplankton community (fig. 9.17). After millions of years of continual deposition, the calcareous foram tests constitute an important component (>75%) of deep oceanic sediments. Other ingredients of deep oceanic sediments are the remains of calcareous holoplankton, such as the calcitic coccoliths (Prymnesiophyta; usually < 10 um in diameter) or aragonite-secreting pteropods (i.e., planktonic gastropods), as well as chitinous remains of copepods (Crustacea), tintinnids (Tintinnidae) and other zooplankton (van Waveren 1989a,b, 1993). We should point out, however, that in other oceanic regions, notably in the higher latitudes as well as in the equatorial east and central Pacific, deep oceanic sediments consist mainly of siliceous (amorphous hydrated silica or opal Si0 2 nH 2 0) remains of diatom (Bacillariophyta) frustules or radiolarian (Radiolaria) skeletons (fig. 9.18). In areas of high productivity, such as the Banda Sea, diatoms can be so productive that their frustules form diatom oozes (Situmorang 1989). In ocean basins where deposition occurs above the lysocline (i.e., depth at which dissolution of calcareous material takes place), foraminiferal tests may form deposits hundreds of metres thick, which are often referred to as "globigerina ooze". In the Banda Sea, foraminiferal ooze is the dominant biogenic deposit, and has been classified according to composition and texture (fig. 9.19). One type consists mostly of tests of pelagic foraminifera Globorotalia spp. and Globigerina spp. and is of a loose sandy character, while a more clayey foram ooze consists of foraminifera in a sticky material (Situmorang 1989). Foraminiferal ooze in fact covers about one-half of the world's deep-sea floor, and thus is the most widespread deposit on the planet. This is especially true along submarine ridges, where depositional environments are generally above the lysocline depth, which varies from sea to sea. Lysocline depth, and thus the fate of foraminiferal tests and other calcareous fragments, is controlled by a number of factors, but mainly by temperature and pressure (Be 1977). Calcium carbonate becomes progressively more soluble with increasing depth (i.e., increase in pressure) and decreasing temperature. However, the dissolution susceptibility of foraminiferan tests is also highly species-specific, since it is also related to their general morphology. The tests of spinose planktonic foraminifera, such as Globigeri-

CONTRIBUTION TO MARINE SEDIMENTS

417 Figure 9.17. Percentage of ocean floor covered by pelagic biogenic sediments.

2.6% Radiolarian ooze

0.6% Pteropod ooze

Figure 9.18. Radiolarian ooze from the Argo Abyssal Plain, depth 5375 m. Leg G.6-2 of the Snellius-ll Expedition. Photo courtesy of M. K. Adisaputra, Bandung.

Pusat Pengembangan

Geologi Kelautan - LIPI,

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FORAMINIFERIDA

Figure 9.19. Globigerina ooze dominated by planktonic foraminifera Globorotalia menardii. Location: Lombok Ridge, depth 3150 m, Leg G.6-4 Snellius-ll Expedition. Photo courtesy of M. K. Adisaputra, Bandung.

Pusat Pengembangan

Geologi Kelautan - LIPI,

noides ruber, G sacculiferand Globigerinella aequilateralis, undergo rapid dissolution as they sink to the bottom, which has a significant effect on the composition of bottom sediments, especially in areas below the lysocline (Troelstra et al. 1989). We should point out, however, that complete dissolution of the foraminiferan tests (and all other calcareous material) takes place mostly on the seabed. More resistant to dissolution is a group of non-spinose foraminifera with thicker tests, such as Neogloboquadrina dutertrei, Pulleniatina obliquiloculata and Globorotalia tumida, which tend to dominate bottom sediments (Troelstra et al. 1989). In the Indonesian Archipelago, the depth of the lysocline in the various basins varies from region to region, depending on the local hydrological conditions. In the Weber Deep, the lysocline is located at a depth of 1700 m, and is characterized by rapid dissolution and fragmentation of foraminiferan tests as well as a significant increase in the abundance of radiolarians. Since radiolarian skeletons are siliceous, they are less prone to dissolution. In addition they may be covered by a thin organic layer providing additional protection. In the Seram Trench, the lysocline is located at a depth of 1100 m, which seems to be one of the shallowest in the archipelago. According to Troelstra et al. (1989), the shallower lysocline in the Seram Trench is associated with strong local upwelling processes, which bring up cooler and calcium-carbonate-undersaturated deep water masses that cause rapid dissolution of foraminiferan tests. The same fate awaits the pteropod (i.e., planktonic

CONTRIBUTION TO MARINE SEDIMENTS

419

mollusc) shells, which are more prone to dissolution, since they consist of pure aragonite. Aragonite is much more soluble than calcite, which is the main constituent of foraminiferan tests. A number of recent studies have been conducted that allow a simple comparison of lysocline depths. Barmawidjaja et al. (1983), working in the northern Molucca Sea, just northwest of Halmahera, found that in an area dominated by the Mindanao Current, the lysocline was located below a depth of 3510 m. In the Sulu Sea, located to the northwest of the Sulawesi Sea, Linsley et al. (1985) reported lysocline at 3800 m, and Berger et al. (1976) reported a lysocline depth of 4100 m in the west Pacific. In the Lombok Basin (Indian Ocean), Adisaputra (1989) positioned the lysocline at a depth of 3157 m, and at 3300 m in the Savu Sea. In the Argo Abyssal Plain south of Nusa Tenggara, the lysocline was at a depth of 4200 m. Thus it seems that the shallowest lysoclines occur in regions of upwelling. In general, deep oceanic water masses are undersaturated with respect to calcite and aragonite. Thus, at a certain depth (>5000 m), most, if not all, carbonate sediments disappear. Oceanographers have coined the termcm-2carbonatecompensation depthcm-2 (CCD) to refer to the depth at which carbonate skeletal material of pelagic organisms contributes less than about 20% of the total sediment. The CCD in the Weber Deep is at a depth of 3300 m (Troelstra et al. 1989). The deepest CCD occurs in the Argo Abyssal Plain at a depth of 5400 m (Adisaputra 1989). As would be expected, the CCD depends on temperature and pressure; however, additional factors such as pelagic and benthic productivity are involved as well. In oceanic systems, sedimentation (i.e., accumulation of sediments) is mainly a function of surface water productivity; thus sedimentation can be used as a proxy indicator of productivity, but with great caution. Sediment traps are frequently employed by oceanographers to determine the downward flux of organic matter. In a study conducted in the Banda Sea, Ganssen et al. (1989) reported sedimentation rates ranging between 11.8 to 15.6 cm.kacm-2. In other parts of the archipelago, sedimentation rates may be much higher, especially in areas influenced by major river systems. Eisma et al. (1989) estimated that based on present-day sedimentation rates of about 60 cm.kacm-21 (c. 150 mg.cm-1.a-1), the Flores Basin will be filled in 10 million years. Based on 8 C values (i.e., -27.24 to -23.72) it seems that most of the sedimenting material is of autochthonous origins. Since Si0 2 is less susceptible to dissolution than carbonate, the glassy skeletons of diatoms and radiolarians are the major component of oceanic sediments below the CCD, and at depths greater than 5000 m. Diatoms also tend to dominate in higher latitudes, or in regions where upwelling promotes high primary productivity (e.g., west coast of Peru), while radiolarian ooze may dominate in low latitudes. Surface productivity of radiolarians, and other planktonic microfossils, is very seasonal in the Banda Sea. Thus, during the non-upwelling period, productivity is very low and surface samples will seldom contain more than a few specimens, yet examination of the deep sediment cores will reveal that productivity during the upwelling period must be considerable. The concurrence of the surface abundance of planktonic microfossils and their benthic assemblages is high in nonupwelling regions year-round, while in areas under strong seasonal upwelling, the concurrence is low during the Northwest Monsoon and high during the Southeast Monsoon. Troelstra et al. (1989) identified over 300 species of radiolaria with some impressive abundance figures. For example, the flanks of the Weber Deep can

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contain as many as 400,000 specimens per gram of CaC0 3-free sediment. These assemblages are dominated by Dictyocorne turncatum, Botryocytris scutum, Phorticium pylonium, Ommatartus tetrathalamus and Lithomellissa thoracites (Troelstra et al. 1989). A high abundance of radiolarians (100,000 to 300,000 specimens per gram of CaC03-free sediment) on the flanks of the Weber Deep has been explained by van de Paverd and Bj0rklund (1989) as being a result of a localized upwellmg that affects primarily the flanks of the Weber Deep. Upwelling on the west flank of the Weber Deep may develop during the Northwest Monsoon as the surface waters are driven to the east, and are being replenished by deeper water along the west flank (van de Paverd and Bj0rklund 1989). The system reverses during the Southeast Monsoon with localized upwelling along the eastern flank of the Weber Deep. However, the Weber Deep has recently been considered not to be influenced by upwelling (Troelstra et al. 1989). As would be expected, higher radiolarian abundance are found in the Seram Trench and on the continental slope of Irian Jaya where upwelling is more pronounced. These assemblages are dominated by Dimelissa apis, Phormacantha histrix, Pseudocobus obeliscus, Lophophaena variabilis and Arachnocorys circumtexta (Troelstra et al. 1989).

Chapter Ten

Environmental Factors INTRODUCTION The Indonesian Archipelago is one of the most complex marine domains on the planet. The complexity of the Indonesian marine environment (i.e., physical, chemical and biological) is a reflection of its geologic history and atmosphericoceanic interactions. Indonesian seas are among the most productive in the world, and support important coastal ancj. offshore fisheries that are a significant source of protein for millions of coastal people throughout the archipelago. The high productivity of the intra-archipelagic seas is related mainly to high discharge rates of terrigenous material (e.g., nutrients and organic matter) from numerous rivers on Sumatra, Java, Kalimantan, Sulawesi and Irian Jaya, and to various types of geostrophic upwellings which are most pronounced in the eastern parts of the archipelago and along the south coast ofJava, Bali and the Lesser Sunda Islands (Wetsteyn et al. 1990; Zijlstra et al. 1990). Significant local upwelling also occurs in numerous straits throughout the archipelago, where high-velocity currents cause vertical entrainment of deeper water masses. In addition, many coral reefs in areas with high tidal ranges generate their own daily upwelling events during flood tides by tidal suction (Wolanski 1992). Cold updrafts experienced by divers along steep reef slopes and drop-offs are a manifestation of this common "upwelling" phenomenon throughout the archipelago. In the eastern Banda Sea, high nutrient concentrations in the 150-300 m layer occur throughout the year, supporting high primary production rates (e.g., 500 g C..m-2.yr-1) (Tijssen et al. 1990; Wetsteyn et al. 1990). In the Aru Basin, primary production rates during the upwelling season (i.e., Southeast Monsoon) were measured at 7-12 g C.m-2.d-1, which are among the highest recorded in tropical seas (Gieskes et al. 1990; Tijssen et al. 1990). A generalization can be made that marine and coastal productivity of the intraarchipelagic seas in the western parts of the archipelago are influenced to a greater degree by terrestrial processes, while the high productivity in the eastern sectors is driven by upwellings. In this chapter we use coral reefs to illustrate how various environmental factors affect shallow-water coastal ecosystems.

TEMPERATURE Temperature and Reefs It was recognized early in the history of coral reef science that coral reef development was limited to tropical and subtropical regions, where surface seawater tem421

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peratures do not get much below 18°C (Dana 1843; Vaughan 1918, 1919; Molengraaff 1928; Yonge 1940). Veron and Minchin (1992) have conclusively demonstrated that functional coral reefs in Japan are restricted to latitudes where sea surface temperatures do not fall below the 18°C minimum for extended periods of time. In addition, they suggested that the development of coral reefs may not be limited by the low temperature tolerance of reef-building corals, but rather, restricted by light, macroalgae and other biotic and abiotic factors (Veron and Minchin 1992). High-latitude reefs in the Indo-Pacific as well as in the Arabian and Atlantic regions puzzled many early scientists. The effects of temperature on the distribution of coral reefs were not always as clear-cut as of other environmental parameters, such as light (Vaughan 1919; Gardiner 1931; Achituv and Dubinsky 1990). However, Darwin (1842) was well aware of the modifying influence of the warm Gulf Stream, which supports extensive reef development, albeit with a depauperate coral diversity, on the Bermuda Platform in the North Atlantic (32° N). The Bermuda Islands, in 32°15' N is the point furthest removed from the equator, in which they [i.e., coral reefs] appear to exist; and it has been suggested, that their extension so far northward in this instance is owing to the warmth of the Gulf Stream.—DARWIN 1842 The northern outpost of Atlantic reef development on the Bermuda Platform is indeed directly attributed to the overriding influence of the warm water masses associated with the Gulf Stream, which originates in the Gulf of Mexico. As a result the sea surface temperatures are not only within the thermal limits of reef-building corals, but are also favourable for biogenic reef accretion. However, seasonal fluctuations in sea surface temperatures do occur, and these are most likely responsible for the lower coral species diversity when compared to the western tropical Atlantic. Recent comparative growth-rate studies of hermatypic corals (colonial and solitary), from Bermuda and the tropical western Atlantic, demonstrated that coral growth rates decrease with increasing latitude, mainly as a function of lower mean sea surface temperatures characteristic of higher latitudes (Tomascik and Logan 1990; Logan and Tomascik 1991). While there is a general consensus that temperature has played a significant role in the Recent (i.e., Holocene) worldwide distribution of coral reefs (Achituv and Dubinsky 1991), the mechanisms through which this influence is manifested remains a topic of numerous research efforts. Direct mortality associated with exposure to low temperature is an obvious mechanism. Early studies on the thermal tolerance of reef-building corals showed that they can tolerate a relatively wide range of temperatures, at least for short time periods (Mayor 1917; Edmonson 1928; Yonge and Nicholls 1931). Thus, some hermatypic species such as Pocillopora damicornis (Indo-Pacific) and Siderastrea radians (Atlantic) are able to survive for a few hours at 15°C (Vaughan and Wells 1943), while other species, such as Astrangia danae (Atlantic), seem to grow well at temperatures as low as 14.5°C (McCloskey 1970). Since reef-building corals seem to have wider thermal limits than coral reefs, thus wider geographic boundaries (both latitudinal and longitudinal), it seems probable that the boundaries of coral reef distribution are linked to subtle temperature effects on coral calcification rates (Wells 1956; Gladfelter et al. 1978;

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423

Jokiel and Coles 1977), growth rates (Glynn and Stewart 1973), and most importantly, to ecological processes that have a direct bearing on the competitive fitness of the reef-builders (i.e., reef constructors) in high-latitude macroalgae-dominated environments (Crossland 1988; Hatcher 1991, 1993). Temperature has a direct effect on the rate of metabolic processes in all organisms, and it has been shown that, with an increase of 10°C, the metabolic activity of various organisms, measured by oxygen consumption, approximately doubles (Hoar 1966). Therefore, it seems likely that the sensitivity of coral species to temperatures below 18°C is inversely proportional to metabolic rates and growth rates. A few coral species can tolerate temperatures as low as 15°C, however, calcification rates are greatly reduced. At temperatures above 31°-32°C, metabolism increases to such rates that calcification and, therefore, growth may also be reduced. Since temperature directly affects the solubility of calcium carbonate (i.e., CaCO3 is more soluble in cold water), it follows that the calcification (i.e., deposition of CaCO3) in reef-building corals may be limited in cold water, restricting vigorous reef development to only tropical and subtropical seas. In these environments biological mediated precipitation of CaCO3 can proceed at a faster rate than physical, chemical and biological destruction (Barnes and Chalker 1990). However, earlier suggestions have recently been questioned that coral reefs were restricted latitudinally because of direct temperature effects on reproduction (Edmondson 1946; Jokiel and Guinther 1978; Kojis and Quinn 1984; Jokiel 1985), survival of coral planulae (Rosen 1975), or feeding rates (Edmondson 1928; van Woesik 1995; Veron 1995). Comparative studies on high-latitude coral reef systems support the hypothesis that temperature plays a key role in the distribution of coral reefs, through direct and/or indirect effects on the structural and functional dynamics of biological communities (Crossland 1982; Wiebe et al. 1982; Johannes et al. 1983). Since coral reefs are products of biogenic processes, the most obvious being calcium carbonate deposition by hermatypic corals and other calcifying organisms, it follows that temperature regulation of calcification rates may be a key mechanism through which temperature effects are manifested in coral reef distribution patterns. The existence of high-latitude coral reefs can be directly attributed to the modifying effects of warm oceanic currents (e.g., Gulf Stream, Kuroshio Current, Leeuwin Current) that create favourable environmental conditions at high latitudes for coral growth (i.e., calcification and skeletal extension). Temperature and Corals The temperature story gets more complicated when we consider the individual coral species tolerance to temperature (Jokiel and Coles 1977), or the tolerance of different regional populations of the same species or genera. It has been stated that corals, and other reef-associated organisms, live at temperatures close to their upper thermal limits (Johannes 1975). It was therefore suggested that because of coral's relatively narrow temperature tolerance range (i.e., stenothermic), temperature increases of only a few degrees above ambient (~ 2°-3°C) can result in reduced growth rates or widespread mortalities in most coral species (Neudecker 1987; Jokiel and Coles 1990). Temperatures exceeding 33°C usually elicit a phenomenon called "coral bleaching", which is the expulsion of the symbiotic zooxanthellae from coral tissue by the coral animal. The presence of a large number of

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bleached (i.e., white in colour) coral colonies is a good indicator of possible environmental stress that may be caused either by natural processes (i.e., increased surface water temperatures during El Nino events) or by human impacts (i.e., thermal power plant effluents or other stressors). It should, however, be pointed out that most of the studies conducted on coral temperature tolerance are based on laboratory experiments. Coles et al. (1976) demonstrated that the upper lethal temperature limits of tropical (i.e., Enewetak Atoll, Marshall Islands) corals are approximately 2°C higher than congeners from subtropical regions (i.e., Hawaii). If these results represent a general pattern, then we would expect corals from regions with higher ambient sea surface temperatures to have a tolerance for higher temperatures. If higher tolerance levels are linked to higher average ambient sea surface temperatures, as was demonstrated by Coles et al. (1976), then this type of adaptation should also occur on regional or even local scales. Field observations indicate that, indeed, some coral species are able to adapt to a relatively wide range of temperatures; thus we may see coral species existing in high-latitude reefs where average temperatures are at the lower limits, but are relatively stable throughout the year. On the opposite end of the spectrum we find corals living in areas with temperatures up to 42°C. Therefore, corals are able to acclimate to a wide range of local temperature regimes within the temperature tolerance range of the species over a short period of time. However, sudden and major deviations from the acclimated temperature regimes may trigger bleaching or be lethal. Figure 10.1 illustrates four different sea surface temperature regimes from four different regions of the Indo-Pacific. The climate in the Marshall Islands and the Moluccas (i.e., Banda Sea) is clearly tropical, while Hawaii is considered by some to be subtropical (Coles et al. 1976). It is clear, however, that average monthly sea surface temperatures in the Marshall Islands do not fluctuate as widely as in the Moluccas. The difference in sea surface temperatures between these two regions is attributed to strong seasonal upwelling in the Banda Sea, which occurs during the Southeast Monsoon (May-September). Note that the upwelling in the Banda and Flores Seas during the Southeast Monsoon also has a considerable cooling effect on the sea surface temperatures in the Java Sea (fig. 10.1). The cooling effect in the Java Sea is probably related to the westward-flowing surface currents which bring in cool upwelled waters from the Banda Sea. The most intense upwelling occurs duringJuly-August. Without the cooling effect, sea surface temperatures in the Java Sea would most likely be higher, probably comparable to sea surface temperatures in Jakarta Bay (fig. 10.2). During the West Monsoon the temperatures in the Banda Sea are considerably higher than in the Marshall Islands. Sea surface temperatures as high as 32°C were measured on a number of occasions on shallow seaward reef slopes in the Banda Islands and southeast Seram (e.g., Koon Island). Across the archipelago in the shallow Java Sea, sea surface temperatures as high as 34°-36°C are common during low tides, especially in shallow lagoons and subtidal reef flats. However, these high temperatures are of relatively short duration, and with the incoming tide sea surface temperatures usually return to normal. Gunung Api (Banda Api), in the Banda Islands, provides evidence that some coral species are able to recruit and survive for long periods in high-temperature environments. Gunung Api is an active volcano with considerable geothermal

TEMPERATURE

425 Figure 10.1. Comparison of average monthly sea surface temperatures in Java Sea, Banda Sea, Hawaii, and Marshall Islands.

Figure 10.2. Annual sea surface temperature fluctuations in Jakarta Bay and Java Sea. Note the cooling effect of the Banda and Flores Seas because of upwelling during the Southeast Monsoon (May to September).

activity, which is indicated by the numerous shallow-water "hot vents" along the coastline. Submarine fissures from the 1988 lava flow on the northwest coast of Gunung Api are the main source of geothermal fluids. Most of the geothermal vents are intertidal; however, one subtidal vent was discovered at a depth of about

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0.5 m below the lowest spring tide. Seawater temperatures within a few metres of the vent were 32°C (ambient 28.7°C), and the thermal surface plume was detected about 75 m down-current. The temperature of the geothermal fluid, a few centimetres above the vent, was 42°C. Growing within 2 m of the vent were numerous healthy colonies ofAcroporaflorida, A. danai, Pocillopora damicornis and P. verrucosa. The temperature recorded was 34°C. Located at the immediate periphery were mature colonies of Pocillopora damicornis, P. verrucosa and Mycedium elephantotus. The temperature at the periphery of the vent was 37°C. The most surprising find (Tomascik and Mah, pers. obs.) was ajuvenile Mycedium elephantotus (4.2 cm in maximum diameter), which was growing inside the vent. Another interesting observation was that the polyps of the juvenile colony were fully extended during daylight, while those of the larger colony, on the periphery of the vent, were not. Seawater temperature in the vent was recorded as 42°C on a number of visits to the site. These observations suggest that scleractinian corals may not be as stenothermic as was once believed. It seems that under natural conditions, scleractinian corals have a high degree of genotypic (i.e., relating to the genetic constitution of an individual or a group) and phenotypic (i.e., relating to the visible properties of an organism or a group) plasticity to tolerate a wide range of environmental conditions. Rapid colonization of the 1988 lava flow on Gunung Api as well as rapid coral skeletal growth rates (but not necessarily calcification rates) may be linked to higher temperatures associated with the geothermal activity (R. Ginsburg, pers. comm.). Coral Reefs and Upwelling As discussed in the preceding section, temperature is now generally accepted to be the main factor determining the global distribution pattern of coral reefs (Veron 1995). Low temperature has been evoked as the main environmental factor responsible for the absence of coral reefs in areas of intense upwelling (Dana 1843), mainly along the western margins of continents. The west coast of South America (i.e., Galapagos Islands) and the west coast of Africa are the two most frequently cited examples. The west coast of Australia is an exception, since it is influenced by the warm Leeuwin Current, which has a modifying influence on the upwelling regime (Hatcher 1991). Considering the tropical nature of the Indonesian Archipelago, and that the archipelagic seas are the major route of the Indonesian Throughflow, which carries the warm western Pacific water (29°C), it seem unlikely that low temperatures are of any concern. Nevertheless, the eastern archipelagic seas are well-known for their intense upwelling during the Southeast Monsoon (Wyrtki 1961), and, therefore, the question of whether this may have an effect on the distribution of coral reefs within the archipelago needs to be addressed. Based on the wide distribution of coral reefs, as well as non-reefal coral communities in the archipelago, it is clear that upwelling per se is not a significant factor determining coral reef distribution. In the eastern part of Indonesia, seasonal upwelling, driven by the Southeast Monsoon, occurs throughout much of the Flores and Banda Seas (Wyrtki 1961), two regions with the highest coral diversity in the archipelago (Best et al. 1989). In contrast, the eastern tropical Pacific, where upwelling systems occur along the western coasts of South and Central America, significantly reduces sea surface temperatures, thus determining regional distribution patterns of coral reefs. However, the seasonal upwelling system in eastern Indonesia

TEMPERATURE

427

does not result in a considerable drop of sea surface temperatures. Sea surface temperatures in upwelling regions are directly related to the intensity of the upwelling, which is mainly a function of wind strength and fetch. It is the intensity of the upwelling that determines the depth from which the water masses originate. The depth of the thermocline is another key factor that determines the sea surface temperatures in upwelling areas. A generalization can be made that sea surface temperatures are much cooler in areas of strong upwelling than in areas of weak upwelling, since the water masses originate from greater depths in the former. Sea surface temperatures in the Banda Sea indicate that the upwelled water masses most likely originate from relatively shallow depths (100-200 m). For example, sea surface water temperatures in the Banda Sea are about 29°-30°C during the Northwest Monsoon (i.e., February/March), and 26°C during the Southeast Monsoon (i.e., August) (Zijlstra and Baars 1987; Boely et al. 1990; Ilahude et al. 1990; Zijlstra et al. 1990). Thus, sea surface temperatures during the strong Southeast Monsoon upwelling period are in fact at the optimal levels for reef development (Jokiel and Coles 1977). It has been shown that seasonal upwelling significantly enhances the primary and secondary productivity of surface waters through nutrient enrichment (Schalk 1987). Optimal temperatures, combined with abundance of plankton, suspended particulate matter and nutrients, may be one explanation why this region contains some of the most diverse coral reef communities in the archipelago. Habitat diversity may be another. Notably it is widely regarded that coral reefs have evolved in oligotrophic oceanic waters, which are often regarded as "nutrient deserts". However, this may be erroneous, since the area in the world that contains the highest coral diversity is also subjected to frequent upwelling and cannot be regarded as oligotrophic; in fact, organisms characteristic of oligotrophic and eutrophic environments are found living side by side (van Woesik, pers. comm.). It must be stressed that high-nutrient conditions that exist during the upwelling period cannot be equated with anthropogenic eutrophication, which has caused severe degradation of coral reefs worldwide (Bell 1992). The seasonal upwelling also results in wider seasonal temperature fluctuations when compared to non-upwelling areas. Nevertheless, narrow sea surface temperature fluctuations are a characteristic feature of the archipelagic seas as illustrated by an example from the Flores and Banda Seas (fig. 10.3). It is clear that while the fluctuations are relatively small (26°-30°C), there are both latitudinal and longitudinal differences. Sea surface temperatures in the Flores Sea do not fluctuate as much as those in the Banda Sea, and the northern latitude stations show narrower ranges than the southern stations at the same longitude. These differences are linked with the more pronounced upwelling events in the Banda Sea as well as along the southern margin of eastern Indonesia. These minor shifts in sea surface temperatures most likely have no effect on coral communities; however, this is still pure speculation since comparative quantitative biological data are lacking. The occurrence of localized upwelling events around individual islands has been corroborated by recent studies in the Banda Islands. In situ temperatures measured over fringing reefs during the upwelling events (range 24°C to 28°C), as well as temperature and nutrient depth profiles taken during the Snellius-II Expedition (Wetsteyn et al. 1990), indicate that the upwelling water masses which regularly spill over the fringing reefs in the Banda Islands originate from depths between 60 to

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100°

105°

110°

115°

120°

125°

130°

135°

140°

Figure 10.3. A) Sea surface temperatures (SST) in Banda Sea, illustrating progressive cooling of surface waters in easterly (left to right) and southerly (top to bottom) directions (i.e., towards the upwelling centres). Weekly data averaged over July 1981 to June 1985. Ill: July-September (Southeast Monsoon and upwelling); IV: October-December (transition); I: January-March (Northwest Monsoon); II: April-June (transition). B) Map of the four areas in Banda and Arafura Seas in A (above). From Boely et a/. 1990. NJSR 25(4) Fig. 21 p.427.

SALINITY

429

80 m. Intense localized upwelling, especially along vertical drop-offs, is also indicated by surface sea conditions (fig. 10.4). Intense tidally-induced upwelling events (current velocities measured at 5-7 knots) may be a significant feature of Indonesian islands, and an important source of nutrients for the reef systems. Nutrient concentrations measured at these depths in the Banda Sea range between 0.13 - 0.53 uM and 0.9 - 7.8 uM of P 0 4 and NO s , respectively (Wetsteyn et al. 1990). High ambient nitrate concentrations may be one explanation for a recently demonstrated 8 N enrichment of Pontes spp. tissue samples collected from a fringing reef on Gunung Api (Heikoop, unpublished data). These values are similar to the 8 N values obtained from corals collected from a sewage-polluted area in the Maldives (Risk et al. 1993). The great diversity of coral reef communities in the Banda Islands suggests that seasonal monsoonallyinduced upwelling, as well as the daily tidally-induced upwelling, may actually promote the productivity of coral reef communities, and thus reef development. High Temperatures High sea surface temperatures are not yet likely to be a significant factor in the regional distribution of coral reefs in the archipelago, however, they may have a significant local effect on community structure and reef zonation. High temperatures have been linked with reef-bleaching events (zooxanthellae expulsion) which resulted in extensive coral mortalities, thus affecting zonation patterns within reef systems. However, because of turbulent mixing associated with strong tidal currents, a dominant feature of the archipelagic seas, high sea surface temperatures, in general, do not seem to be a major factor in reef distribution patterns. The 1983 ENSO-associated mass-bleaching event in the Java Sea has been attributed to a significant temperature rise; however, it occurred in a relatively isolated area and under a specific set of environmental conditions. In addition, all impacted reefs during the 1983 ENSO event were in areas of intensive development and under heavy exploitation pressure. For example, coral reefs in Kepulauan Seribu are being subjected to a variety of anthropogenic impacts (e.g., eutrophication, siltation, etc.) which may have made them more susceptible to natural disturbances. In other regions of the archipelago, temperatures as high as 36°C have been observed to cause no visible damage to corals (fig. 10.5), and healthy corals have been found to survive periodic exposures to temperatures as high as 42°C (fig. 10.6). Based on these observations, it may be concluded that the general coral reef distribution patterns within the archipelago are not temperature-related.

SALINITY

Definitions In the simplest terms, salinity is the measure of the sum of all inorganic salts dissolved in seawater. The average concentration of dissolved salts in the world's oceans is roughly 3.5% by weight (i.e., 35 psu - physical salinity units, which is equivalent to ppt and %o) (Dietrich et al. 1980). Salinity in the oceans varies greatly

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Figure 10.4. A) Tidallyinduced upwelling eddy along a reef d r o p - o f f at Koon Island, southeast Seram (eddy temperature 23°C; reefal sea surface temperature 28.5°C). B) Current-induced upwelling eddy in Boleng Strait, between Adonara, Solor and Lembata Islands, East Nusa Tenggara (eddy temperature 22°C; normal sea surface temperature 29.4°C). Photos by Tomas and Anmarie Tomascik.

among regions as well as with depth. Salinity is a conservative property of seawater, since changing the concentration of major salts (i.e., dissolved ions such as Cf and Na+, etc.) will not alter their ratios (i.e., relative proportions), which remain constant irrespective of salinity changes. With the exception of the Red Sea and the Arabian region, oceanic sea surface salinities are highest in the Atlantic Ocean. Salinity changes have a profound effect on the physical properties of seawater, since salinity causes variations in specific gravity, and therefore density, which has a direct effect on the movement of water masses and circulation patterns. Within the Indonesian Archipelago, sea surface salinities vary greatly along both spatial and temporal scales. Different climatic conditions throughout the archipelago, large islands with high volumes of river runoff, various geological barriers (i.e., island arcs), and sea surface circulation patterns all contribute to significant salinity differences among the various archipelagic seas. It is well-known

SALINITY

431 Figure 10.5. Sangalaki Island, Berau Islands, East Kalimantan. The large shelf reef with a densely vegetated cay is located about 60 km offshore. Sea surface temperatures on the intertidal reef flat reach 36°C at low tides (tidal range 2.75 m). Intrusions of upwelled water masses are common. Photo by Tomas and Anmarie Tomascik.

Figure 10.6. Banda Api shallow-water coral community subjected to hydrothermal fluids and daily heating at low tide (tidal range 2.5 m). Sea surface temperatures as high as 38°C were recorded, however, coral bleaching is not a major problem. Photo by Tomas and Anmarie Tomascik.

that chemical properties of seawater have a direct effect on the basic metabolism (osmoregulation and other physiological processes) of marine organisms, including corals, and therefore, salinity has a significant influence on their distribution, including the distribution of coral reefs. For many reef organisms, and especially corals, salinity determines survival, spawning areas, optimum growth conditions, and movements during life history changes. In the following discussion on the effects of salinity, most experimental studies to date were on mature coral colonies. Recent experiments in Guam and Okinawa have clearly demonstrated that a 20% reduction in seawater salinity causes an 86% reduction in fertilization success in corals (Richmond 1993). The question here is whether regional and local salinity differences, within the archipelagic seas, are sufficient to affect the distribution of coral reef communities and coral reef development in general.

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Salinity and Global Reef Distribution The fact that coral reefs were absent from all tropical regions heavily influenced by river runoff is a well-known fact since Darwin (1842). Salinity, at first, seems a simple explanation, since extensive coral reef development has never been reported in any region where the average sea surface salinities are below approximately 29 psu. This includes all regions affected by the world's major river systems (table 10.1). For example, in the Atlantic, major coral development occurs in regions where sea surface salinities range between 34-36 psu (i.e., Caribbean). The Amazon River, with an annual runoff of about 5.7 x 1012 m3, has a far-reaching influence on the distribution of coral reefs along the northwest coast of Brazil, and perhaps also influences the general nutrient dynamics of the eastern Caribbean Sea. As a result of the North Equatorial Current, the Amazon low-salinity plume travels in a northwesterly direction, effectively inhibiting major reef development until the Lesser Antilles (i.e., Trinidad and Tobago), about 2100 km away. These low-salinity eddies, which are distinct water masses, are seasonally detectable as far north as Barbados, 2400 km from the Amazon delta. The northwest coastline of Brazil is, however, fringed by extensive seagrass beds and mangal forests, while the Brazilian Abrolhos to the south support flourishing coral reef communities. Recent questions concerning the regional effects of the Amazon River plume on Caribbean reef nutrient dynamics has attracted new research efforts in the region. Coral reef regions of the Pacific Ocean and the central Indian Ocean (e.g., Maldives) maybe considered as relatively unaffected by continental processes (with the exception of some high oceanic islands such as Hawaii, Kosrae, etc.). Sea surface salinities within the subtropical and tropical regions of the Pacific Ocean range between 35-36 psu, with the exception of a small region extending from about 90° W to 130° W (between 10° N20° N), where the average salinities are about 34 psu (fig. 10.7) In the Southeast Asian region, coral reefs occupy a variety of environments,

Table 10.1. Mean annual flow rates of the Mahakam and Solo Rivers, and some of the major world river systems. River

Continent

Amazon Congo Orinoco Mississippi St. Lawrence Danube Rhine Yangtze Mekong Brahmaputra Mahakam Solo

South America Africa South America North America North America Europe Europe Asia Asia Asia East Kalimantan East Java

Source: Dietrich et al. 1980; Dutrieux 1991; Hoekstra

1980.

Average annual runoff (103m3.sec-1) 180.0 42.0 28.0 17.5 10.4 6.5 2.2 34.0 15.9 20.0 1.5 0.8

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433

Figure 10.7. Distribution of sea surface salinities throughout the world's oceans. From Open University

1991.

ranging from the turbid shallow continental seas (e.g., Gulf of Thailand, Java Sea, Arafura Sea) to deep oceanic basins where atolls and high-island reefs (i.e., fringing and barrier type) developed in clear oceanic waters free of terrestrial influence (e.g. Andaman Sea, Sulawesi Sea, Flores Sea, Banda Sea). The western extent of the Southeast Asian coral reef distribution can be delineated by the Andaman and Nicobar Islands, located to the north-northwest of Sumatra. The Andaman-Nicobar reefs developed on a submarine ridge that was formed at the northern extension of the Sunda Trench subduction system. In fact, the Andaman-Nicobar Ridge is a continuation of the non-volcanic Mantawai-Nias Outer Arc Ridge, which is part of a continuous subduction system along the Sumatra Trench (Hutchison 1989). These reef systems are sufficiently distant from the Ganges-Brahmaputra and Irrawaddy River systems, whose combined annual runoff volume is about 1.1 x 10 m . As would be expected, the Ganges-Brahmaputra delta (the largest subaerial delta in the world) supports extensive mangal forests. In the Gulf of Thailand, river runoff has a significant effect on the distribution of coral reefs. Low sea surface salinities (20 psu) in the northern part of the gulf can extend up to 70-80 km seaward from the Chao Phraya River delta (Piyakarnchana 1981), thus limiting coral development to the more southern regions of the gulf. Upon close inspection, however, the salinity control hypothesis becomes problematic, especially in the Southeast Asian region. River runoff not only dilutes the salinity of the receiving coastal waters, thus subjecting marine organisms to osmotic stress, but all major tropical river systems carry a heavy load of suspended sediments that may be transported long distances from the river, particularly during storm

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events. Thus, a clear separation between salinity, turbidity and sedimentation as the major causative factors in reef distribution is extremely difficult to demonstrate (Fagerstrom 1987). Interestingly enough, Darwin (1842) was of the opinion that, while reduced salinities unquestionably have a negative impact on corals (i.e., physiological stress), the main factor responsible for determining the geomorphology and distribution of inshore coral reefs, at least on a local scale, is the deposition of fluvial sediments (i.e., sedimentation). This process seems to be the main factor determining the distribution of many inshore reefs within the Indonesian Archipelago and Southeast Asia in general. No doubt brackish water would prevent or retard the growth of coral; but I believe that the mud and sand, which is deposited, even by rivulets when flooded, is a much more efficient check.—DARWIN 1842 In some arid regions of the world (e.g., Bahrain), reef-building coral communities are known to flourish in environments where salinities seldom drop below 42 psu. Extensive reef development in the Red Sea and the Arabian region supports this fact. For example, Pontes nodifera-dommaXed communities form patch reefs in environments where salinities range between 43-45 psu (Sheppard 1988). Shortterm exposure to high salinities have been reported earlier, but Sheppard (1988) found that some coral species (e.g., Porites nodifera, Siderastrea savignyana and Cyphastrea microphthalma) can survive in salinities as high as 50 psu. However, it is evident that only marginal coral reef development occurs in these high-salinity environments, and the coral communities are dominated by only a few tolerant coral species (28 species in Bahrain) (Sheppard 1988). Based on the wealth of information from the Red Sea and Arabian regions, it seems that scleractinian corals have a remarkable physiological tolerance (i.e., acclimatization ability) to high saline conditions (Coles 1993). However, coral survival or reef development at salinities similar to or below the normal minimum have not been demonstrated (i.e., 15 psu below the lower normal value of 30 psu). Vaughan (1919) found that some Caribbean coral species survived in salinities of 28 psu for 24 hours, while other studies indicated greater sensitivity. Note that coral tolerance to low salinities, under laboratory conditions, is not indicative of their ability to survive under these conditions as viable populations, nor is it indicative of reef-forming capability. What the laboratory experiments have shown is that corals are able to tolerate above- or below-normal salinities, for certain periods, which is not surprising since on a global scale coral reefs encounter a wide range of salinities (Stoddartl969). Coles and Jokiel (1992) have reviewed the available salinity data from numerous coral reef studies around the world, and came to the conclusion that: "corals and other reef organisms can live in normal salinities as low as 25 %o and as high as 45 %o". However, out of the 38 studies cited in the review, only six studies reported salinities below 30 psu, the lowest value of 25 psu being from the Gulf of Thailand. Closer inspection of the Thailand data (Piyakarnchana 1981) reveals that the reefs in the northern regions of the gulf were outside the 25 psu isohaline. The fact that low salinity water is usually restricted to the 1 to 3-m-deep surface layer was not considered. The reevaluation of the available data (Coles and Jokiel 1992) suggests that there is no evidence to support the view that vigorous coral reef development occurs at salinities below about 29 psu. It is also clear that reef-building corals are

SALINITY

435

not as stenohaline as was once assumed. In fact, many coral species have a wider salinity tolerance range than other reef-associated organisms. Based on the worldwide distribution of coral reefs, and experimental data, it seems that reef-building corals are able to tolerate salinities of about 25 psu for brief periods. This suggests that periodic exposures to low salinities, which may, for example, occur during tropical storms, will not inhibit reef development. It may be concluded that reef-building corals and coral-reef-associated organisms generally (there are exceptions) flourish in regions where average sea surface salinities are maintained between 30 to 36 psu, and therefore, coral-dominated "reef-building" may be loosely defined as a stenohaline process. Salinity and Indonesian Reef Distribution Sea surface salinities in the archipelagic seas are influenced by a number of environmental factors (e.g., runoff, precipitation, evaporation, surface current patterns) . The most obvious, and the most easily interpretable, is the effect of river runoff during the Northwest Monsoon. Heavy rainfall during the wet season is often associated with floods and significant fluctuations in river discharge volumes. For example, during the Northwest Monsoon (Novemberjanuary) the daily flow rates of the Solo River, with a catchment area of about 16,000 km2, may fluctuate between 300 and 1800 m3.sec4, while during the Southeast Monsoon the daily flow rates fluctuate between 80-250 m3.sec"1 (Hoekstra 1989; Hoekstra et al. 1989). Just to the south of the Solo River delta, the Brantas River discharges an additional 600-1200 m .sec" of runoff during the wet season (Heokstra et al. 1989), while the combined volume of these two major Javanese rivers (maximum wet season combined flow of 3000 m3.sec"1) is considerably less when compared to the flow rates of some of the other major rivers in Southeast Asia (e.g., 15,900 m .sec" for the Mekong River). These rivers nevertheless have a significant impact on the coastal geomorphology and the distribution of shallow-water coastal ecosystems along the northeast coast ofJava and Madura. Since the coastal surface currents during the Northwest Monsoon are dominated by a net eastward flow, the Solo-Brantas River runoff may be a key factor responsible for the absence of major reef development along the north coast of Madura. Significant reef development does not occur until about 130 km to the east of the Solo-Brantas deltas, along the south coast of Madura (e.g., Pulau Giligilingan). However, marginal fringing reefs occur along the south coast of Madura, about 40 km east of the Solo-Brantas deltas (e.g., Labuan). The nearest reefs to the Solo delta in the east Java Sea are the fringing and patch reefs around Pulau Bawen, about 130 km to the north. However, it is suggested that the general absence of coral reefs in this region of the east Java Sea is probably associated mainly with the lack of suitable substrate, rather than with the effects of the SoloBrantas River runoff. This is an interesting area of research in an accessible area. As expected, there seems to be a general absence of coral reefs throughout the archipelago in the immediate vicinity of major rivers. However, closer inspection reveals that the effects of low salinities may be confounded by other environmental factors, such as reduced light levels associated with increased turbidity or sedimentation, as well as possible interactions of these parameters. Heavy runoff from the major rivers on Java Island (e.g., Solo River) during the Northwest Monsoon is most likely the main factor responsible for the absence of reefs along much of its

436

ENVIRONMENTAL FACTORS

northern coastline. To what extent salinity influences this outcome is difficult to determine, since clear-water, low-salinity, tropical coastal environments have not been studied. Lack of suitable substrate, high turbidity and/or heavy sedimentation along much of the northern coastline may be the overriding factors causing a dearth of reef development. There are a number of patch reefs situated off the north coast ofJava (e.g., Gosong Tengah, Karang Rakit, Gosong Pamanukan, etc.) where the presence of suitable substrate has allowed coral reef development. These areas were originally far enough from the influence of natural river runoff, and as a result, coral communities became established and reefs were formed. Hydrological conditions along much of Java's north coast have changed in recent years (Verstappen 1953), mainly as a result of major changes in land drainage patterns. These changes in coastal hydrodynamics are associated with large-scale conversion of coastal plains into agricultural land (i.e., wetland paddy fields) as well as deforestation and land conversion of upper water catchment areas. Both of these practices have resulted in increased freshwater runoff and a high load of suspended solids carried down the rivers during the wet monsoon. As a result of macro-scale weather patterns associated with the monsoonal system, as well as the high mountain ranges, the Indonesian Archipelago is not climatically homogeneous. When viewed from a regional level, there is a distinct annual rainfall pattern present. Figure 10.8 illustrates that the west (e.g., Sumatra, Kalimantan, Java) and east (e.g., Irian Jaya) regions of the archipelago receive considerably more rainfall on an annual basis than the central regions (e.g., Lesser Sunda Islands, Moluccas). The lowest annual rainfall occurs in the Lesser Sunda Islands. The central region of the archipelago, with the Philippines, was a known centre of generic coral diversity (Best et al. 1989; Veron 1986, 1995). However, Hoeksema (see box 7.2; chapter 7) recently expanded the generic diversity of this region; thus for the time being it may be considered as a high-diversity sub-centre within the Indonesian-Philippine Province. The differences in rainfall between the various regions of the archipelago, however, are not only of quantity. With regards to coral reef development and coral community structure, the seasonal patterns of rainfall may be more significant than the average annual values. It is often stated that the archipelago is under the influence of a monsoonal climate, and therefore, marked seasonality in rainfall should be expected. However, as is clearly demonstrated in figure 10.9, there are distinct differences not only in the amount of rainfall, but also in the seasonality. For example, average monthly rainfall in Fak-Fak, located on the Onin Peninsula in Irian Jaya, ranges between 250-400 mm without any pronounced seasonal fluctuations. The seasonal weather pattern in this area can best be described as either wet or very wet. Similar conditions occur in Tarakan, East Kalimantan, where the average monthly rainfall varies between 250-300 mm throughout the year. Strong dry-wet seasonality is, however, characteristic of much of the archipelago, especially in the southeastern regions (i.e., Flores, Timor, Komodo, etc.). Smaller areas within each region can also exhibit marked differences as a result of microclimate variability. In these areas the climate is influenced mainly by local physiography. Local differences in seasonal patterns in average monthly rainfall are illustrated in figure 10.10. These regional differences in rainfall patterns have a measurable effect on the hydrology of the coastal area, and therefore, on the distribution of coral reefs. The coastiine of most eastern islands is fringed by extensive

SALINITY

437 Figure 10.8. Average yearly rainfall (mm) for the main physiographic regions of the Indonesian Archipelago. Source: RePPProT 1990.

Figure 10.9. Average monthly rainfall (mm) for Kalimantan, Java, Nusa Tenggara and Irian Jaya, demonstrating different climatic conditions. Continually wet climate - Irian Jaya; Slight seasonality in rainfall - Kalimantan; Dry and wet monsoons Java; Dry and arid conditions - Lesser Sunda Islands. Source: RePPProT 1990.

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ENVIRONMENTAL FACTORS

Figure 10.10. Local rainfall patterns. Comparison of average monthly rainfall between Larantuka, east Flores and Buyasuri, Lembata Island. The two districts belong to the East Flores Regency, East Nusa Tenggara. The distance between the two locations is 60-70 km. Rainfall data from 1990. Source: Pemerintah Kabupaten DATI II, Flores Timur 1992.

fringing reefs. For example, the island of Yamdena, in the Tanimbar Islands, has a fringing reef that runs virtually uninterrupted along the entire length (i.e., c. 375 km) of the coastline, except where breached by rivers and rivulets. Some of the most diverse coral communities are found in the drier eastern regions of the archipelago, characterized by strong seasonality in rainfall and low volume of land runoff. However, this is not to say that diverse coral communities, or extensive coral reef systems, have not developed in areas of continually high rainfall. The extensive barrier reef system in the Berau Islands, in East Kalimantan, illustrates that major coral reef development can occur in relatively turbid, shallow-water coastal areas in close vicinity to major river systems (see fig. 10.5). Flow rates in the Berau River may peak at 500 m .sec" following heavy rains, and carry a considerable amount of suspended solids. During December SPM concentrations in the delta may reach 175 mg.l"1, however, available data indicate that average SPM concentrations in the Berau River are about 80 mg.l" (Hatfindo, pers. comm.). The fluvial fan of the Berau delta extends some 20 km offshore. To the north of the Berau Islands, however, the coastal water quality rapidly deteriorates from the influence of the Bulungan and Sesayap Rivers. Karang Baliktaba is the last coral outpost at the northern fringe of the Berau Barrier Reef (fig. 10.11). The Baliktaba Reef is located about 30 km to the southeast of the Bulungan

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439

Figure 10.11. The Baliktaba reef complex is the last coral outpost at the northern boundary of the Berau Barrier Reef (see fig. 12.4, chapter.12). The reef rises from a depth of about 50 m and is surrounded by a mixture of fluvial and bioclastic sediments. A) The intertidal and shallow subtidal habitats are dominated by acroporids, supporting a diverse reef fish fauna. B) A more diverse coral community is found on the deeper reef slope. The reef extends to a depth of about 15-20 m. Photos by Tomas and Anmarie

Tomascik.

River delta, just at the outer fringes of the turbidity zone associated with the Bulungan River plume. The coastline between Karang Baliktaba and Tarakan is fringed by extensive mangrove and Nypa fruticans forests. However, high turbidity of the coastal waters loaded with fluvial sediments seems to be the main factor responsible for the absence of coral reefs from this area, since surface salinities are above 30 psu. The lack of a suitable substrate also may be important, considering that the shallow-water habitats seaward of the mangrove belt are mainly mud flats. Shallow-water and high-velocity tidal currents rapidly mix the river water with coastal water masses, and help to maintain high concentrations of fluvial sediments in suspension for long distances. The fluvial fan of the Bulungan River extends over 20 km offshore, where extensive mud banks have been formed. The complexity of the abiotic factors makes it impossible to ascertain what role, if any, salinity plays in the distribution of coral reefs in this region. Effects of Low Salinities Sea surface salinity fluctuations, associated with river runoff and/or heavy rainfall, may be a key factor determining local (i.e., within reef) zonation patterns, especially in coastal areas with large tidal ranges. Coral reefs located in close proximity to large rivers (e.g., Berau River, East Kalimantan) exhibit a marked transition from low-diversity reef flat communities which are dominated by a few hardy coral species (e.g., Pontes lobata, P. lutea, P. cylindrica, Montipora turtlensis, M. mollis, M. hispida, Goniastrea favulus, G. aspera, Goniopora djiboutiensis, G stokesi, G. lobata) that are generally tolerant of high turbidity and lower salinities, to high-diversity reef slopes that support luxuriant coral growth further from the river (fig. 10.12). In contrast, reef flats of oceanic platform reefs (e.g., Nil Desperandun, Banda Sea) are

440

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Figure 10.12. The Tanjung Batu Reef is located about 1.5 km offshore, at the seaward limit of the Berau River delta (see fig. 12.4 for location). The apparently lagoonless shelf reef, with an area of about 30 km2, is cut by numerous narrow channels. A) Mainly as a result of its sheltered location, the intertidal reef flat coral community, on Tanjung Batu Reef, has a distinct lagoonal character (tidal range 3 m). The Berau River plume frequently impinges on the reef; however, as a result of strong southward currents the reef is maintained relatively sediment-free. Note the presence of Pontes microatolls. B) The deeper reef slope (4-5 m) of Tanjung Batu Reef has noticeably higher coral diversity and coral cover, corresponding to increased water clarity. Photos by Tomas and Anmarie

Tomascik.

high-energy coral communities, dominated mainly by luxuriant assemblages of acroporids (e.g., Acroporapalifera, A. robusta, A. danai, A. austera, A. cytherea, etc.) as well as other coral groups (fig. 10.13). Note, however, that oceanographic conditions (i.e., currents and waves) determine zonation patterns of oceanic coral reef communities removed from the influence of terrestrial processes. Based on observations from East Kalimantan, the north coast ofJava, and the southwest coast of Irian Jaya (e.g., Bintuni Bay, etc.) we arrive at the expected conclusion that coral reefs are not found along coastlines characterized by predominantly estuarine conditions (e.g., widely fluctuating salinities and high turbidity), and extensive tidal range. Unfortunately, the role of salinity in this wellknown dictum is not very clear. However, it is well recognized that substantial damage and mortality can occur with flood waters that significantly reduce sea surface salinities below ambient. For example, Glazebrook and van Woesik (1993) and van Woesik et al. (1995) documented that corals suffered about 85% mortality when floodwater reduced salinities to 8%o following a major storm. Dead corals were rapidly overgrown by turf algae two weeks after the flood event. Coral mortality was restricted to depths less than 1.3 m, however, reef-bleaching was observed at greater depths. Tissue samples collected from corals subjected to low-salinity stress indicated hypertrophy, hyperplasia and lysis of the epidermis as well as degenerative changes in the endodermis, which sometimes extended to necrosis (Glazebrook and van Woesik 1993; van Woesik et al. 1995). In addition, bacterial emboli

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Figure 10.13. Subtidal reef flat on Nil Desperandun, a submerged oceanic platform reef in the Banda Sea. The subtidal reef flat is dominated by a diverse Acropora community, predominantly by Acropora palifera and A. robusta. Tidal currents sweeping over the reef were recorded at about 4.5 m.sec"1 (8.7 knots). Photo by Tomas and Anmarie

Tomascik.

were present in the subepidermal tissue layer (van Woesik et al. 1995). Sea surface salinities as low as 25 psu (found in the Gulf of Thailand) may not be uncommon in many coral reef areas, and have indeed been observed in Jakarta Bay during the Northwest Monsoon (Tomascik, pers. obs.) as well as during the regular oceanographic cruises of the R.V. Samudra. Ilahude and Liasaputra (1980) reviewed oceanographic reports from Jakarta Bay and their results clearly demonstrate a sea surface salinity gradient extending from the Citarum and Bekasi Rivers in a northwesterly direction towards Kepulauan Seribu (fig. 10.14). While sea surface salinity data are useful in helping to delineate the horizontal boundaries of river plumes, depth profiles should accompany the salinity data to obtain the complete picture. Table 10.2 demonstrates that low sea surface salinities do not necessarily extend into the lower water column (i.e., freshwater, being less dense than seawater, floats on top), and therefore, may not directly impact upon deeper coral reef communities. Similar observations were made by Kastoro et al. (1989) just offshore of the Solo and Brantas River deltas in East Java, during the wet and dry seasons. The combined discharge volume of both rivers is considerable during the wet season, but the salinity effect is rather localized due to rapid mixing with seawater. During the rainy season, sea surface salinities 1 km offshore (depth 0.3-1 m) ranged between 25-28 psu. At 2 km offshore (depth 1-2 m) salinities

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Figure 10.14. Normal sea surface salinities in Jakarta Bay during the Northwest Monsoon (January), characterized by heavy rainfall and river runoff. From llahude and Liasaputra

1980.

Table 10.2. Salinity depth profile for Stations 44 and 45 in Natuna Sea, Oceanographical Cruise Report No. 3, R.V. Samudra, February 12 - March 5, 1971. Profiles at Stations No. 44 and 45 taken on February 25, 1971. No. 44: 02°14'00" N, 110°18'00" E\ No. 45: 02°32'40" N, 109°56'10" E. Depth

Salinity

(psu)

Jm) 0 10 20 30 45 50

No. 44 28.22 32.81 32.86 33.01 33.03

No. 45 28.35 33.19 33.22 33.22 33.22 -

Temperature

(°C)_

No. 44 27.4 25.99 25.99 25.88 25.83

No. 45 27.5 25.82 25.85 25.85 25.81 -

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increase to 30-31 psu, and 4 km offshore (depths 3.5-9 m) salinities ranged between 30-33 psu. Note, that the coral reefs of Jakarta Bay were not long ago flourishing systems, which apparently evolved in relatively close proximity to major rivers (Verwey 1930a; Umbgrove 1939a). Low sea surface salinities such as those found in Jakarta , Bay require careful interpretation, since low-salinity water masses seem to be restricted mainly to the surface layer and generally do not extend below 3 m depth. This seems to be the case for Jakarta Bay, where a low-salinity surface layer (24.5 psu) was observed to extend to a depth of 2 m below the surface, thus not in direct contact with the depauperate coral community 5 m below. However, all rivers running into the Java Sea are heavily laden with sediments, and therefore, it is more than likely that significant reduction of light levels and sedimentation are the main impacts associated with low-salinity surface water masses. The effects of light, turbidity and sedimentation are discussed in the next section. Berau Islands: Case Study The influence of river runoff, and the difficulty in determining the effects of salinity, can be illustrated by an example from Karang Buliulin, which is located at the southern extension of the Berau Barrier Reef system, about 40 km downstream of the Berau River delta (see fig. 12.4). The intertidal reef flat, about 35 km in area, is under the influence of a strong diurnal tidal regime (F=0.23), with a maximum tidal range of about 2.8 m (fig 10.15). The reef flat has strong 'lagoonal' attributes in terms of sediment characteristics and community structure. There are visually striking differences in terms of coral community structure and coral cover between the western and the eastern reef slopes (fig. 10.16). This difference may be attributed to the influence of the Berau River, whose sediment-laden plume impacts first on the west slope of the reef. The sediments at the bottom of the reef slope, along the west boundary of the reef, contain a higher fluvial fraction when compared to the more bioclastic sediments found along the eastern boundary of the reef. In essence, the extensive intertidal reef flat of Buliulin Reef (4 km wide and 11 km long) acts as a "block", since water quality (in terms of turbidity) between the west and east regions of the reef are very different. Horizontal underwater visibility (at a depth of 5 m) on the west slopes was 5 m, while on the east coast the horizontal visibility extended to about 15 m. This area offers interesting research opportunities. Salinity measurements demonstrated that localized upwelling of deeper water masses during the flood tide modify the effect of low-salinity eddies originating from the Berau River plume. At the end of the Northwest Monsoon, a large low alinity eddy (28 psu at 0.5 m) was detected over the entire reef area (Tomascik and Mah, pers. obs.). At high spring tide, the reef flat is about 2.5 m below sea level. Salinity at the bottom was 33 psu, thus the surface low-salinity plume was restricted to the top 1.5 m depth. During the flood tide, the tidal currents are flowing in a predominantly northwesterly direction with a velocity of about 1 m.sec . Sea surface temperatures measured over the reef flat (about 50 m from the outer reef slope) averaged about 30°C, while the bottom (2.5 m depth) reef flat temperatures and salinities were 28°C and 33 psu, respectively. The sharp temperature and salinity differences between the surface and bottom layers is related to localized upwelling

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Figure 10.15. An example of a semidiurnal tidal cycle (F = 0.23) over a two-day period in the Berau Islands. Maximum amplitude is about 2.8 m. Source: Daftar Pasang Surut 1994.

Figure 10.16, A) Reef slope on the west side of Karang Buliulin, Berau Islands, East Kalimantan. The west side of the large offshore patch reef is influenced by river runoff. B) Reef slope on the east side of Karang Buliulin. Note higher diversity, cover and improved water quality. Photos A and B were taken on the same day, within one hour. Photos by Tomas and Anmarie

Tomascik.

SALINITY

445 Figure 10.17. Possible model of localized upwelling induced by high-velocity tidal currents (tidally-induced upwelling on Buliulin Reef). A) Upwelling by tidal suction during flood tide. B) Vertical entrainment by tidal jet during ebb flow. After Wolanski 1992.

generated by flood-tide entrainment of deeper water masses along the seaward edge of the reef (maximum depth 60 m) (fig. 10.17), known as the Bernouilli effect or tidal suction (Thompson and Golding 1981; Thompson and Wolanski 1984). The upwelling of more saline and cooler waters provides an effective barrier against the influence of low-salinity surface water masses, especially if this phenomenon is cyclic (i.e., tidal). However, the flourishing coral communities on both sides of the reef do not extend below the 10 m isobath, where there is an abrupt transition from hard substrate to fine carbonate sediments mixed with fluvial deposits. Thus, the depth extent of the coral communities seems to be substrate-limited rather than water quality or light-related. Based on 10-minute manta-tow surveys around the periphery of the reef, it became apparent that the southeast sector of the reef had a significantly patchier distribution of coral communities, when compared to the north, east and west sectors. The patchiness at the southeast sector of the Buliulin Reef was attributed to numerous shallow channels, which were the main outlets for the reef flat water during the ebb flow (fig. 10.18). The location of the channels corresponds well with the predominant net southward flow in the region. Based on these limited observations, we may nonetheless speculate that low salinity is not the main limiting factor in this region of the archipelago. Turbidity and sedimentation associated with river runoff, or with highly productive reef flats, are most probably the dominant environmental factors determining the regional and local distribution of coral communities and coral reefs.

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ENVIRONMENTAL FACTORS

Figure 10.18. A) Sediment channel at the southeast sector of Buliulin Reef at high tide (depth 3 m). B). Sediment accumulation occurs at the base of the reef slope (depth 10 m; ebb flow). Note the greatly reduced water clarity.

Effects of High Salinities Sea surface salinities are also influenced by evaporation which is controlled by solar radiation, by the regional and local wind fields, and by humidity. During the dry season, weather conditions may result in rapid heating and evaporation of extensive shallow intertidal pools and shallow lagoons, thus substantially increasing sea surface salinities. Under these conditions the formation of high-temperature (>31°C) and-salinity (>36 psu) water in the shallow reef flat pools may have a significant influence on the lower reef slope coral communities through a cascading mechanism (i.e., high-density water will spill over the reef slope). Downwelling of warm, saline, sediment-laden, reef flat water masses is a common occurrence, especially on reefs with extensive reef flats and steep dropoffs. For example, during a survey of an offshore reef on the northeast coast of Koon Island (southeast Seram), strong local downwelling was documented along the reef wall at a depth of 10 m. The circulation in this area is dominated by strong tidal currents that can reach velocities of up to 3 m.sec'. The downwelling was characterized by a sharp change in temperature, from 28.2°C to 30.1°C, and a strong downward current with a velocity of about 0.5 m.sec"1. The turbid water was restricted to within 20 m of the reef wall, and provided a sharp contrast to the off-reef oceanic waters. The downwelling cell extended for about 300 m along the reef wall, where it stopped as abruptly as it had begun. Large feeding aggregations of planktivorous fish (e.g., Chromis amboinensis, C. atripes, C. xanthochir, C. analis, Pseudanthias luzonensis, P. dispar, P. lori) were oriented in an up-reef position, against the downwelling, in sharp contrast to the horizontal orientation of large schools of Pterocaesio randalli feeding about 15 m off the reef wall. During intense ENSO events, these normal downwelling episodes may be greatly intensified in certain regions (e.g., Java Sea), triggering reef-bleaching events, which may be followed by substantial coral mortality (Brown and Suharsono 1990). Furthermore, the downwelling may be an important mechanism for the

L I G H T A N D C O R A L REEFS

447

transport of fine-grained, shallow-water reefal sediments into deeper parts of the reef slope. Sediment cascading down the reef slope was observed on numerous occasions on most oceanic platform reefs in the Banda Sea.

LIGHT AND CORAL R E E F S The ability of corals to build reefs using the energy of the sun, is the key to the existence of all modern coral reefs, and perhaps all reefs in all geological time—VERON 1995

The statement certainly rings true, considering the fact that the coral-dinoflagellate holobiont is an autotrophic biological unit, whose complex biochemical processes are clearly light-dependent (Lewis 1974c; Jaubert 1977; Gattuso andjaubert 1984; Kinzie 1987; Titlyanov 1987; Waymann et al. 1987; Falkowski et al. 1990; Harland et al. 1992). Perhaps with the exception of some temperate species (e.g., Astrangia danae), coral-zooxanthellae symbiosis is an obligate association, since all zooxanthellae corals rely to a great extent on their algal symbionts for nutrition (i.e., reefbuilding corals can derive up to 95% of all nutritional requirements from the zooxanthellae), and, most importantly, calcification (Porter 1976; Crossland et al. 1980b; Muscatine et al. 1981; Davies 1984; Muscatine 1990). It is, therefore, not surprising that solar energy is one of the key environmental factors limiting the distribution of zooxanthellate corals, and coral reefs in general. Sunlight is the primary force that drives coral reef ecosystems, as well as all other autotrophic systems on the planet. In fact, with the exception of the chemosynthetic (i.e., chemical energy) deep oceanic benthic communities associated with hydro thermal vents, the earth's biosphere is driven by the Sun's energy. In simple terms, the basic fuel for the coral-zooxanthellae holobiont is the Sun's electromagnetic radiation which can be used by the zooxanthellae. The basic aspects of light that will have an effect on the coral symbionts is light quantity (the amount of energy delivered) and light quality (spectral composition), both of which are affected by a variety of factors. Both quantity and quality of light are greatly influenced by seasonal changes in day length (i.e., amount of energy received) and the angle of incident sunlight (i.e., quality of energy received). Therefore both the amount and spectral quality of insolation will decrease with increasing latitude (fig. 10.19). However, after reviewing the available literature, it is apparent that the amount of solar radiation reaching the coral communities at their northern and southern limits of distribution is clearly not a limiting factor, since corals have been observed to grow at depths of 30-40 m at latitudes 31° S and 35° N in clear-water .environments (Veron 1995). When we look at the distribution of the total amount of solar energy being received at various regions on Earth, it becomes apparent that the expected latitudinal gradient is not as clearly defined as one would expect (fig. 10.20). There are obvious differences between the Arctic and equatorial regions, but at lower latitudes (50° N - 50° S) the pattern is not clearly apparent. For example, Indonesia, straddling the equator, is expected to receive the highest doses of solar

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ENVIRONMENTAL FACTORS

Figure 10.19. Average amounts of incident solar radiation over one year p e r i o d , expressed as gm.cal.cm"2day"1 at latitudes 0°, 30° N and 52° N. Source: Raymond 1963.

Figure 10.20. Average amount of solar radiation (W.rrf2) that is received at the earth's surface. From Open University 1991b. Open. Univ. Ocean. Circ; p. 159; fig. 6.1.

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449

energy, but in fact receives less over the year than the Red Sea, which is situated at the northern limit of coral reef distribution in the Indian Ocean. The average yearly amount of solar energy that is being received in Indonesia, ranges from 155 to 230 W.m 2 , well within the range being received by the coral communities situated at the southern limit of their Pacific distribution at Lord Howe Island (31° S). Yet these systems differ greatly in terms of reef development and coral species richness. The differences in the amount of solar energy being received at any one place on the planet, are closely linked with atmospheric conditions, which have a significant effect on insolation (i.e., the actual amount of solar energy reaching Earth's surface). As a general rule, large continental areas, on average, receive more insolation than oceanic regions, especially in the equatorial belt. The major factor responsible for lower insolation over equatorial oceans, as well as over tropical rain forests, is rapid evaporation which results in cloud formation. A higher concentration of water vapour in the atmosphere, and formation of clouds over most equatorial regions significantly increase light scattering, absorption and reflection, which ultimately results in reduced insolation. These processes are fundamental to global, regional and local energy distribution patterns along both spatial and temporal scales. On any given clear sunny day in the tropical latitudes, the sea surface around noon receives, on average, 2500 uE.m^.sec"1 of photosynthetically active (available) radiation (PAR), which can potentially be utilized in algal photosynthesis. However, not all of this surface energy is available to marine autotrophs, since the moment light reaches the sea surface it is quantitatively and qualitatively altered by numerous physical processes. Only a small fraction of the available solar energy reaches the autotrophs, and an even smaller fraction is actually used in photosynthesis. The absorbed fraction is usually referred to as the photosynthetically usable radiation (PUR), and for zooxanthellae, and other marine algae, it is restricted mainly to radiant energy between 400 to 700 nm (i.e., nm = 10" m). The amount of PUR actually used in photosynthesis varies among marine autotrophs, and depends entirely on the composition of their photosynthetic pigments (e.g., chlorophyll), which determines the wavelength-specific absorption spectrum of the photosynthetic cells (Falkowski et al. 1990). Zooxanthellae have two absorption peaks within the PAR range. A wide absorption peak occurs between 400 and 550 nm, while a second narrower peak occurs between 650 and 700 nm . While solar radiation does not seem to be a limiting factor in the latitudinal distribution of coral reefs, it is the key factor that determines their depth distribution (Bak and Luckhurst 1980; Sheppard 1982; Huston 1985). Considering the fact that the majority of zooxanthellate corals are restricted to the euphotic zone, it may be said that the key limiting factor with regards to depth distribution is symbiosis itself. Even under optimum conditions (i.e., clear oceanic water), zooxanthellate corals are generally restricted to depths between the surface and 100 m (Wells 1957, 1969), however, the most vigorous coral reef development seems to occur at depths of 3 to 20 m. In most tropical seas (e.g., Flores Sea), the 18°C isotherm lies roughly at depths between 150-200 m, which is below the euphotic zone even in the clearest ocean. Temperature is therefore not a major factor limiting depth distribution of corals in most tropical seas; however, it may be a factor in some upwelling

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regions (e.g., Galapagos). On the other hand, symbiosis is most likely the key factor that has allowed the zooxanthellate corals to widely disperse throughout the oligotrophic tropical oceanic regions, although their dependence on symbiosis seems to severely limit their competitive ability, and possibly wider distribution, in high-latitude macroalgaedominated environments (Crossland 1982,1988; Hatcher 1991,1993; Veron 1995). The role of symbiosis in environments with an abundant food supply (i.e., upwelling areas characterized by a high abundance of zooplankton) has not been investigated. The nature of the symbiotic relationship is discussed in chapter 7. Light Environment in Indonesia Coral reefs within the Indonesian Archipelago are situated in an equatorial region, and therefore, the question of whether light has an effect on the horizontal distribution patterns within the archipelagic seas may seem to be a non-issue. However, as was discussed earlier, the archipelago is not climatically homogeneous. In fact, there are significant climatic differences between the western (i.e., Sumatra, Java, Kalimantan and most of North and Central Sulawesi), and eastern (i.e., east of Java, except Irian Jaya) regions of the archipelago. One of the most obvious differences is the average annual rainfall, which is loosely indicative of the amount of insolation. The southeastern regions, on average, receive considerably higher insolation (between 200-230 W.m"2) than the western regions (150-190 W.m"2). These differences may be related to the drier, at times arid, climate of the Lesser Sunda Islands and southern Moluccas, a wide region influenced by the Australian continental mass to the south. In contrast, the mountain systems and extensive rain forests of Kalimantan, Sumatra, and Malaysia have a combined regional influence on cloud formation in the area, thus the climate is slightly more humid and cooler than in the eastern regions. Whether there is an ecologically meaningful correlation between insolation and coral reef distribution patterns within the archipelago remains to be seen, but it is not likely. The actual amount of insolation actually reaching the various seas of the archipelago is not the main issue, since as far as coral light requirements are concerned, all the regions receive an excess of solar energy year-round. What is, however, important, is how much of the available insolation actually reaches the autotrophic communities, whether they are benthic or planktonic. The amount and quality of available light that reaches the autotrophs in the sea, depends on the synergistic effects of a number of environmental factors, and not just on the amount of insolation. Much of the incident light at the sea surface is reflected back into the atmosphere as a result of surface irregularities caused by waves (i.e., a function of the wind) and the angle of the sun. Even in the clearest of ocean water, 65% of irradiance that actually penetrates the sea surface is lost within the first metre of the water column, with only 1% reaching 60 m, and at about 200 m, less than 0.1% remains. Under ideal oceanic conditions, light penetration is primarily controlled by the optical properties of seawater, and by water column productivity. This certainly is the case in highly productive oceanic systems, such as the Banda Sea, where light penetration (i.e., as measured by a Secchi disk) has a distinct seasonal pattern that is primarily related to seasonal fluctuations in primary and secondary production rates (fig. 10.21). The attenuation of light (through absorption and scat-

LIGHT AND CORAL REEFS

451 Figure 10.21. Correlation between water transparency as measured by Secchi disc (m), and primary production as measured by chlorophyll-a concentrations, Banda Sea. Source: Gieskes et al. 1989.

tering) as it passes through the water column in coastal waters is influenced by a variety of other factors that will be discussed in the following section. While the differences in insolation among the various regions of the archipelago are relatively minor, and most likely biologically and ecologically insignificant, as far as reef and coral distributions are concerned, the amount of solar energy reaching benthic communities clearly varies among regions. Higher rainfall and different geological setting in the western regions of the archipelago, and Irian Jaya, are two key factors (we shall ignore anthropogenic factors for now) that may have a significant influence on both the quantity and quality of land runoff that enters the coastal waters. In terms of coral distribution patterns, land runoff is of primary importance, since the suspended particulate load which it carries, directly influences light penetration and ultimately the amount of light reaching the benthic autotrophic communities. Turbidity and Suspended Particulate Matter. D E F I N I T I O N S . Turbidity can simply be defined as water clarity or transparency. However, while being governed by the optical properties of seawater, turbidity is mainly a function of suspended particulate matter (SPM) concentrations and dis-

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Figure 10.22. Comparison of water clarity (means ± STD) between the Northwest Monsoon, a period of downwelling, and the Eastern Monsoon, a period of upwelling, in the Banda Sea. Lower water clarity is primarily a function of high phytoplankton and zooplankton biomass during upwelling (May-September). Conversely, high water transparency is directly related to low primary productivity during the Northwest Monsoon (November-April). Source: Gieskes et al. 1989.

solved organic compounds in the water column. There is an obvious positive relationship between turbidity and SPM, but the synergistic effects are complicated. Turbidity mainly relates to light-dependent processes, while SPM, in addition to directly affecting turbidity, has a number of other effects that should be treated separately. Clear separation between the effects of turbidity and SPM are difficult, and rarely studied in corals (Yamazoto 1986). Suspended particulate matter can be either biotic (e.g., phytoplankton, zooplankton, organic detritus, etc.) or inorganic (e.g., silt, sand, etc.) in composition. In oceanic environments, removed from the influence of continental processes, turbidity is generally low (i.e., low SPM concentrations; 0.5 mm in diameter) in the water column, and are commonly referred to as marine snow (Shanks and Trent 1979; Alldredge and Silver 1988). Associated with SPM and marine snow are various dissolved organic substances which are either metabolic wastes (e.g., fecal pellets), or products of organic decomposition. The synergistic effect of SPM and dissolved substances in the water column plays a significant role in light extinction. While high concentrations of SPM, and therefore turbidity, may have negative impacts on autotrophs, the nutrient-rich marine snow (comprising most of SPM) is an important source of food for pelagic fauna (e.g., zooplankton and ichthyoplankton) (Shanks and Trent 1979; Alldredge and Silver 1988). As oceanic water passes over the reef, SPM concentrations increase as a direct result of increased productivity (benthic and pelagic), particularly in lagoonal waters where relatively turbid conditions predominate. Oceanic and reefal detrital SPM is an important source of food for benthic filter and suspension-(particle) feeding organisms, and perhaps planktivorous reef fish as well. Furthermore, much of this drifting organic matter can be utilized directly by corals themselves (Lewis and Price 1975; Sorokin 1978). While being an autotroph [photosynthesis-to-respiration ratio (P/R) >1], the coral-zooxanthellae holobiont is also a highly efficient predator, equipped with a variety of auxiliaryfeeding habits (Muscatine and Porter 1977), one of which is particulate or suspension-feeding (Sorokin 1973; Lewis and Price 1975; Lewis 1977b). Earlier studies in the Banda Sea (Nontji 1975) indicated that seasonal fluctuations in primary production, and therefore water clarity, were monsoon-driven

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and associated with upwelling (Wyrtki 1961). Lowest transparency occurs during the Southeast Monsoon (May-September), when strong easterly winds set up a large-scale upwelling system that significantly increases phytoplankton and zooplankton biomass in the water column. During the Northwest Monsoon, which is a period of general downwelling (i.e., sinking) (Wyrtki 1961), the water clarity improves considerably as both phytoplankton and zooplankton biomass drop. At the peak of upwelling, high primary and secondary production in the water column considerably reduces water clarity and thus the depth of the euphotic zone. While higher productivity may be beneficial for coral communities in general, especially at depths of less than 20 m, significant reduction in water transparency during the upwelling period may effectively reduce the maximum depth limit of reef-building. In the productive coastal waters, light penetration becomes a significant factor in both the horizontal (regional to local) and vertical distribution of coral communities. Coastal Environments. In comparison, most coastal waters in the archipelago are characterized by high turbidity, and the suspended particulate matter is of various biotic and abiotic origins, both aquatic and terrestrial. However, in addition to the SPM, coastal waters are known to contain various soluble pigments or 'yellow substances' that are responsible for significant absorption of light (Jerlov 1951). In tropical coastal environments, the bulk of the dissolved organic acids and other substances originate from mangrove forests and tidal swamplands. The influence of terrestrial runoff on inshore coral reef communities has been clearly demonstrated in some areas, by the detection of humic and fluvic acids in the skeletons of massive species of Pontes (Isdale 1984; Boto and Isdale 1985; Susie et al. 1991). On the Great Barrier Reef, humic acids have been detected as far as 80 km offshore, with highest concentrations being recorded in the inshore coastal waters. The main source of humic and fluvic acids is terrestrial plant matter, and according to Susie et al. (1991), soils of tropical rain forests are particularly rich in these organic compounds. Humic acids are continually being leached into the soil and to small, streams that eventually discharge into the coastal waters. However, during heavy rainfall, tropical rain forests lose a considerable amount of surface soil material (contrary to popular belief) that may be particularly rich in these substances. The work by P. Isdale and his colleagues at the Australian Institute of Marine Science (AIMS) clearly demonstrated that corals (i.e., Pontes spp.) are able to incorporate humic acids into their skeletons during calcification. Since humic acid concentrations in most tropical coastal waters are related to the annual patterns of rainfall (i.e., wet-dry), and river runoff, the skeletons of massive corals-when exposed to UV light-exhibit distinct yellow-green fluorescent bands. The periodicity of the fluorescent bands is closely correlated with the amount and frequency of rainfall and the intensity (i.e., volume) of river discharge (Susie et al. 1991). Humic acid concentrations in the coastal waters are, on average, lower during the dry season than during the wet season. During the wet season, high concentrations frequently occur following periods of heavy rainfall. This discovery has helped the Australian researchers to reconstruct past river discharge volumes, and to discern patterns of anthropogenic influences in the main drainage areas bordering the Great Barrier Reef. In contrast, comparable success with fluorescent banding in Indonesia has

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not materialized (Scoffin 1986). However, it should be noted that fluorescent banding studies have been restricted mainly to coral reef areas along the north coast of Java (i.e.,Kepulauan Seribu andjepara). In all studies conducted so far, clear bright UV banding is visible in all coral skeletons, but seasonal patterns are not discernible (Scoffin 1986). UV-banding patterns in corals from Kepulauan Seribu surprisingly show no significant correlation with mainland rainfall data, as they do on the Great Barrier Reef. At first this may be surprising, considering the fact that most parts of the archipelago, includingJava, are under the influence of the monsoonal climate with distinct dry and wet seasons. Seasonal periodicity in the volume of river discharge has been documented for some rivers (e.g., Solo and Brantas Rivers in East Java), and therefore, one would expect to find strong fluorescent banding patterns in most inshore corals. Closer examination reveals, however, that the main study sites are located in coastal areas severely impacted upon by anthropogenic influences. Of significance are the greatly altered coastal drainage patterns and the deforestation (agricultural conversion) of the tropical forests in source watersheds. Seasonal UV banding in corals along the north coast of Java, which may well have existed under pristine environmental conditions in the distant past, is confounded by anthropogenic influences that have greatly modified previous natural environmental conditions along much of the coastline. If UV fluorescent banding in massive corals is indeed closely linked to land runoff, as seems to be the case in Australia, then the lack of banding in Kepulauan Seribu corals may indicate major changes in coastal environment have occurred, and corals are now subjected to the effects of continental processes at higher intensities and frequencies. Effects on Coral Reef Distribution. It has often been stated, and is a widely held view in popular literature, that coral reefs have evolved in clear and oligotrophic (i.e., low nutrient availability) oceanic waters. An often-stated analogy is that a coral reef living in clear nutrient-poor tropical oceanic waters is like an oasis in the desert. This maybe true, but only for the 'blue-water' oceanic reefs (i.e., atolls) of the Pacific and Indian Oceans. Unfortunately, the direct corollary of this assertion is that vigorous coral reef development occurs mainly under the above 'presumed' ideal oceanic conditions, which implies that coral reefs in the more turbid coastal regions are living under less-than-optimal conditions. From the resource management and conservation perspective, this is a potentially dangerous view, since, at least from the Indonesian experience, these turbid, yet highly productive, areas are often ignored and excluded from management initiatives. Indeed, conservation efforts in Indonesia are too often, if not exclusively, directed at the traditionally 'blue-water reefs' (i.e., high tourism potential) of the eastern regions. As was clearly pointed out by Potts (1983), about 50% of all Indo-Pacific reefs are found along shallow continental shelf margins. An unknown percentage of these reefs are scattered throughout the shallow continental seas of the Indonesian Archipelago (e.g., Java Sea, Natuna Sea, Arafura Sea), and along still relatively undocumented coastlines of Kalimantan, Irian Jaya and other large islands. Water clarity along the coastlines of these large islands is considerably less when compared to the oceanic waters of the Banda and Flores Seas, where light penetration is high. For example, the extinction coefficient (k) along the northwest coast of Java ranges between 0.85 m" and 1.7 m", while in the west Java Sea &'ranges

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457

between 0.34 m" to 0.09 m". In comparison, &'in the Banda Sea varies between 0.09 m" during the upwelling season, and 0.05 m~ during the Northwest Monsoon when upwelling is inhibited by predominantly westerly winds. It is unfortunate , that, by focusing most research on the 'classical' blue-water reefs of the IndoPacific (i.e., atolls), a few generalizations (e.g., low nutrients, high water clarity) have permeated the popular literature, even though widely read papers have provided solid evidence that coral reefs have a remarkable capacity to develop and 'flourish' under relatively turbid conditions (Verwey 1931a). But we see that even in the tropics, so close to the equator, and even in a coral reef region, we need not always have the blue clear water one dreams of in thinking about reef formation.—VERWEY 1931A Turbidity is, however, considered to be more important than salinity in inhibiting reef development, even in regions where other environmental factors, such as temperature and substrate, are favourable (Guilcher 1988). As was pointed out earlier, the coral-reef-depauperate coastlines of Southeast Asia, India and the northwest coast of Brazil are influenced by turbidity plumes from major river systems, which confound or override the effects of salinity. In Indonesia, there is a distinct longitudinal gradient of water clarity which may have some influence on reef distribution and development (fig. 10.25). Comparative coral reef studies along this longitudinal gradient have not been conducted, and the results of individual reef surveys in the various regions are not comparable because of varying methodologies and reporting. While some coral data are now available, a common deficiency is the lack of an environmental component. Thus, while we may discern significant differences between different regions in terms of coral community structure (i.e., diversity or coral cover), the main question as to why these differences occur remains unanswered. Nonetheless, the effects of turbidity on the coral community structure and function are wellknown and have been documented elsewhere (fig. 10.26). In many coastal areas of the Indonesian Archipelago coral reefs flourish or persist, albeit within a very narrow depth range, under turbid conditions (e.g., &=1.0). As an extreme example let's examine the coral reefs of Pulau Panjang which are located in Banten Bay (West Java). Coastal waters in Banten Bay are muddy brown most of the time, with conditions improving as one moves further offshore. Turbidity in the bay is high and a Secchi disk disappears in less than 2 m. The entire Banten Bay is influenced by heavy land runoff from intensively used agricultural lands, which has considerably increased soil erosion, and therefore, sedimentation rates in the bay area (Kiswara 1992). However, all islands in the bay are fringed by reefs. The intertidal reef flats are dominated by seagrass communities consisting mainly of Cymodocea serrulata, Halophila ovalis, H. uninervis, and Thalassia hemprichii, which, however, are under considerable siltation stress (Kiswara 1992). The seaward edge of the reef flat (1.5-2 m below sea level) is the only zone on the reef where corals are able to survive without visible trauma, and where visible recruitment was observed (i.e., abundance of small, 700 mm.day , posing considerable safety hazards to large structures such as bridges (RePPProT 1990). While some storms can generate almost cyclone-force winds, the duration and fetch are insufficient for the establishment of heavy sea conditions, which rarely exceed 3 m amplitude. Most severe storms seem to occur in the shallow Java Sea, which has a relatively long fetch-length along an east-to-west axis. The influence of the steady Southeast Monsoon winds on the geomorphology of the Kepulauan Seribu Reefs is well-known (Molengraaff 1928; Kuenen 1933; Umbgrove 1948). Because of the seasonality and relative predictability of the monsoons, coral communities within the archipelagic seas exposed to monsoonal winds have developed characteristic zonation patterns that differ greatly from those in sheltered environments. However, the geological complexity, on regional and local scales, offers a variety of exposed and sheltered shallow-water habitats that modify the influence of atmospheric disturbances. The highly diverse coral reefs (including fringing reefs, barrier reefs and atolls) of the Togian Islands, in the Gulf of Tomini, are a good example of a sheltered system. The Paleleh Mountains to the north (i.e., the North Arm of Sulawesi) and the Balinggara and Batui Mountains to the south (i.e., the East Arm of Sulawesi) provide an effective protection against the winds of the Northwest and Southeast Monsoons respectively (fig. 11.2).

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Box. 11.1. Local weather phenomena over the Indonesian Archipelago. P. A. Winarso, Meteorological and Geophysical Agency of Indonesia, Jakarta, Indonesia. The Indonesian Archipelago occupies a large equatorial region between 06° N and 12° S and 90° E to 150° E. Land represents only about 33% of the total area of the archipelago, thus Indonesia is a maritime nation. The Indonesian Archipelago is under the influence of the Maritime Continental Air Mass, resulting in a humid and warm climate. Because of the characteristics of this air mass and convective activity, clouds are easily generated over this region, causing frequent local storm activity. The effects of local storm events are mainly due to increased turbulence associated with strong gusts of wind. Each year, localized storms, associated with the formation of cumulonimbus clouds, are responsible for environmental damage to shallow-water coastal and marine communities as well as to land. On average, local storms on land and sea are responsible for more damage to the environment and property than any other natural phenomenon (i.e., earthquakes or volcanic eruptions). Local storms may cause storm surges and high seas that can severely impact upon coastal structures, causing severe damage. Heavy seas associated with strong gusts of wind may also impact on benthic communities in coastal areas as well as deeper water offshore. Some of the most damaging storms are associated with the formation of highly convective cells known as the cumulonimbus clouds. The cumulonimbus clouds can form rapidly over open seas, with the base of the cloud usually a few hundred metres above sea level, while the top of the clouds can reach altitudes of 15,000 m above sea level. Wind velocities associated with these events, also known as squalls, can reach gale force for brief periods, causing considerable damage, mainly above water. Because of their short duration they do not seem to have a significant impact on coral communities in general, however, exceptions may occur. These storms are most frequent between November and April, and are closely linked to the position of the Inter-Tropical Convergence Zone (ITCZ). In the ITCZ, air masses from the Northern and Southern Hemispheres converge as a result of high surface temperatures. The ITCZ is characterized by frequent storms with heavy rainfall and strong winds. In addition to the winds associated with the convective cloud activity, strong winds also predominate during the tropical cyclone season of both hemispheres. During these periods local storm surges occur in Indonesian regions that are close to the disturbances. For example, during the tropical cyclone season in the Northern Hemisphere (i.e., May - October), several coastal regions in North Sumatra, East Kalimantan and North Sulawesi may experience storm surges associated with passing cyclones. During the Southern Hemisphere cyclone season, between November - April, coastal areas of Java, Bali, Nusa Tenggara, South Sulawesi and East Timor may experience strong winds associated from the peripheral influence of tropical cyclones in the Southern Hemisphere. The western regions of the archipelago are also influenced by periodic cold surges originating to the north of the South China Sea. The cold surges occur from November to March, with surface winds that can reach up to 60 km.hr"1, creating heavy sea conditions in the Natuna and Java Seas. Heavy seas associated with the cold surges cause considerable damage along the southeast coast of Sumatra, the Riau Islands and the north coast of Java. The storm conditions can be intensified by the presence of the ITCZ, resulting in heavy rainfall and strong, gale-force winds over the shallow Java Sea.

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The reefs around the western Togian Islands are situated in a zone, famous for the unruffled calm and undisturbed quiet of the sea... Neither on the barrier reefs, nor on the atolls, nor on the fringing reefs of the Togian Islands does one single shingle rampart occur. In vain I searched for Montipora foliosa,. . .— • UMBGROVE 1947

The protection from the Northwesterly and Southeasterly Monsoonal winds offered by the two mountain ranges has a pronounced effect on the overall geomorphology of Togian reefs. This is especially so in the northwest region of the island chain, since the islands themselves (elevation >500 m) provide additional protection from the influence of strong Southeast Monsoon winds. The most noticeable effect is on the reef zonation as well as on coral community composition. For example, Montipora foliosa was recognized by Umbgrove (1947) as a conspicuous component on the exposed reef slope of Nyamuk Besar (Leiden), a coral cay in Jakarta Bay. Coral communities on Nyamuk Besar are under strong monsoonal influence. M. foliosa was especially abundant along the northwest and southeast slopes of the shingle rampart, demonstrating clear preference for high-energy environments. The northwest and southeast slopes of the reef face the prevailing Northwest and Southeast Monsoon winds, respectively (fig. 11.3). M. foliosa is also an abundant species (encrusting morphology) throughout the Kepulauan Seribu (i.e., Thousand Islands), especially on the shallow exposed reef slopes. Suharsono (1992) found that M. foliosa was one of the 30 coral species present in all of his sampling stations at Pulau Genteng, Thousand Islands, out of a total of 180 scleractinians. Yet this species is apparentiy absent from the protected shallow reef slopes of the Togian reefs, even though it is found in deeper water (Umbgrove 1947). In contrast, the delicate hispidose Acropora turaki (\yallace 1994) was a dominant species along the southern coast of Pulau Talatakoh, with extensive lagoonal environments protected by a barrier reef. In fact, the calm waters of the Togian Islands seem to be an optimal environment for delicate Acropora species (e.g., Acropora jacquelineae [Wallace 1994]; A. lokani [Wallace 1994]; A. caroliniana Nemenzo) which dominated the shallow reefs as well as offshore atoll environments (C. Wallace and J. Wolstenholme, pers. comm.). In contrast to both Kepulauan Seribu and, the Togian Islands, the oceanic reefs in the Banda Sea (e.g., Nil Desperandun) under oceanic conditions have a much different community. The shallow reef flat and reef slopes are dominated by extensive and dense stands of Acropora palifera, A. cuneata, A. brueggemanni, A. monticulosa and especially A. robusta. The presence of this particular assemblage of acroporids tells us that Nil Desperandun is a high-energy environment subjected to strong wave action and currents. Indeed, during our survey of the reef, current velocity was in excess of 2 m.sec", unusually strong for an oceanic platform reef with a tidal amplitude of about 1.5 m. This superficial look at three different environments suggests that atmospheric disturbances other than tropical cyclones play an insignificant role in the structuring of the shallow-water coral communities in the Togian Islands. Furthermore, the lack of shingle ramparts or coral cays in the region suggests that wind and waves may have negligible roles in the geomorphology of the reef structures. Clearly, this unique region, with an amazing diversity of reef habitats and geomorphologies, deserves a lot more attention from the scientific community than it has received

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Figure 11.3. Wind roses for the central and eastern archipelago during January (A) and July (B). Peta Cuaca Perairan Indonesia

1992.

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493

Figure 11.4. Frequency of tropical cyclones in the Indo-Pacific region. Modified from Scoff in 1993.

thus far. The coral communities in Kepulauan Seribu, and in the past, in Jakarta Bay as well, show a classical zonation structure that developed in response to predictable monsoonal weather patterns. The coral community on Nil Desperandun indicates that large-scale disturbances may occur. Tropical Cyclones On the numerous reefs of the Great Barrier Reef, one can see large coral boulders that were thrown upon the reefs by rough seas during numerous tropical cyclones. Their noticeable absence from the Indonesian reefs that are nonetheless heavily influenced by strong monsoonal winds has been noted previously (Kuenen 1933; Umbgrove 1947). Based on their absence, one can arrive at the conclusion that tropical cyclones do not occur in the archipelagic seas. This generalization does actually apply to most of the vast archipelag;, however, as was mentioned earlier, the southern islands of East Nusa Tenggara, and the northern islands from Biak to Sangihe, have been known to be affected by cyclonic conditions resulting from a bypassing cyclone. Figure 11.4 gives us some idea on the frequency of cyclones that pass in close proximity to the northern and southern regions of the archipelago. Figures 11.1 and 11.4 clearly demonstrate that the Indonesian Archipelago lies between two main tropical cyclone regions. However, because of Indonesia's geographic position and prevailing atmospheric conditions, most tropical cyclones move away from the archipelago. Periodically, however, cyclones do wander in, and with rather devastating impacts to the coral communities. Figure 11.5 shows the general tracks of cyclones that have, in the past, strayed into the eastern regions of the archipelago.

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Figure 11.5. Tracks of cyclones in East Nusa Tenggara and Moluccas from 1908 to 1974. From RePPProT 1990.

Figure 11.6. Tropical cyclone tracts on the Great Barrier Reef and Western Australia from 1908 to 1981. Courtesy of T. Done and the Australian Institute of Marine Science,

Townsville.

CYCLONE "LENA" (JANUARY 2 3 , 1993)

495

There seems to be a well-defined "cyclone corridor" around Timor. Some reports suggest that cyclonic conditions have been recorded as far west as Lombok (RePPProT 1990). The impact of these storms has not been documented previously, and therefore, very little is known about their effect on the coral reef community. If the cyclones are of sufficient strength, we would expect some coral boulders to be thrown upon the reef flats of, at least, the most exposed reefs. However, as was also noted by Umbgrove (1947), corroborative evidence from the islands around Timor is not available. In contrast to their relatively infrequent visits to the eastern part of the archipelago, cyclones are a yearly occurrence on the Great Barrier Reef (fig. 11.6).

CYCLONE " L E N A " (JANUARY 23,

1993)

Maumere Bay is renowned not only for the world-class diving, but also for an amazing diversity of coral reef fish. Because of the high diversity of physiographic features found in Maumere Bay, reef diversity is correspondingly high (i.e., fringing reefs, steep drop-off reefs, a barrier reef, and an offshore atoll). As a result of high habitat diversity, the reefs of Maumere Bay support diverse assemblages of corals. Unfortunately, little work has been done on coral community structure in this area, and therefore little information is available on their diversity. However, the numerous reef habitats support an amazing diversity of reef fish fauna, with over 1700 species identified. (G. Allen, pers. comm.). Figure 11.5 illustrates that tropical cyclones are a relatively rare occurrence in the eastern regions of the archipelagic seas. The last cyclone to pass through the region was "Sally", in December 1971. According to the weather reports, the eye of Cyclone "Lena" did not actually enter the Indonesian seas, however, Flores and the adjacent islands to the east were affected by the winds at the outer periphery of the disturbance. Winds blowing predominantly from the north-northwest (330-350°) reached gusts of up to 45 knots (Beaufort 9 or 80 km.hr 1 ) (Wai-Oti Meteorological Station, Maumere). Considerable structural damage (e.g., shingles and slate removed from roofs) was reported for properties along the coastline. Unfortunately for many of the shallow-water coral communities, the storm occurred within two months of a major earthquake, which caused severe structural damage to many coral reefs in the area. The extent of cyclone damage in the affected regions is not known, but an example from Maumere Bay may serve to illustrate the destructive power of these disturbances, especially in regions where they rarely occur. Figure 11.9 summarizes our qualitative observations on the impact of Cyclone "Lena" on coral communities at various locations in Maumere Bay. Consistent with the direction of prevailing winds, the least damage to the coral communities occurred on reefs located in relatively sheltered areas along the southeast coast of Pulau Besar and Pulau Dambilah. The most severe damage occurred along the south shore of Maumere Bay, which was fully exposed to the storm. However, the area was also struck by the December 1992 tsunami. Which disturbance caused more damage to coral communities is difficult to ascertain. Observations of well-sorted beach deposits indicate that the storm disturbance may have had a greater impact than the 2-4 metre-high tsunami. Rough weather conditions

496

CORAL REEFS: NATURAL DISTURBANCES

Box 11.2. Tropical Cyclone "Lena", January 1993. P. A. Winarso; Meteorological and Geophysical Agency of Indonesia, Jakarta, Indonesia. During January - February 1993, a large-scale atmospheric disturbance occurred over most of the Indonesian Archipelago, extending from the South China Sea through the Java and Flores Seas, to the Arafura Sea. This large-scale atmospheric disturbance was characterized by exceptionally rough seas throughout this region, resulting in the sinking of a number of ships, damage to several oil platforms in the Java Sea, as well as damage to shallow-water coral communities from Java to Flores. Detailed analysis of daily weather reports indicate that two low pressure systems were established in the region, one just northwest of Australia and one over the Gulf of Carpentaria (figs. 11.7 and 11.8). The northwestern depression strengthened into Cyclone "Lena", which moved with a trochoidal-motion track, initially in a westerly direction. "Lena" was, however, deflected from the normal southwesterly direction eastwards. During this time, the area was also experiencing cold surges that are active in January. Associated with the cold surges were strong southerly winds of up to 30 knots (56 km.hr"1) over the South China Sea, which generated fast southerly surface currents of more than 3 m.s"1 (10.8 km.hr"1), and waves of up to 3 m in amplitude. Similar conditions were observed in the west Java Sea (Bangka Island), where 1 d (3 m.s" ), 10-m-deep surface currents and waves of more than 4 m height were recorded. Associated with this atmospheric disturbance were heavy rains of up to 100 mm along most of the north coast of Java and the islands of Nusa Tenggara (i.e., from Bali to Timor).

Figure 11.7. Surface weather analyses on 25 January 1993 at 00.00 UTC.

CYCLONE "LENA" (JANUARY 2 3 , 1 9 9 3 )

497

Box 11.2. (Continued.) Review of historical weather data indicates that these conditions are infrequent, and thus this event may be classified as a rare occurrence, which poses a question as to what may have caused it. It is suggested that the disturbance was a result of the synergistic effect of the abnormal track of Cyclone "Lena" that coincided with the activity of the cold surge from the South China Sea waters. Normally, tropical cyclones in this region move in a southwesterly direction, and the winds that are generated along the southern part of the Indonesian Archipelago are less than 30 knots. Furthermore, this disturbance followed the "El Nino - Southern Oscillation" (ENSO) event of 1991-1992, which incidentally coincided with the peak sunspot activity (11 -year cycle), characterized by increasing activity of solar flares. It is postulated that if post-ENSO periods coincide with increased surface warming due to increased intensity of solar flares, these types of atmospheric events may occur at infrequent intervals. It is expected that meteorological and oceanographic conditions associated with tropical cyclones will continue to affect some parts of the Indonesian Archipelago from time to time. Historical review of large-scale weather patterns and oceanographic conditions may allow us to make future predictions on the possibility of occurrence of these rare events in the archipelago.

Figure 11.8. Satellite cloud photograph of Cyclone "Lena" on 25 January 1993 at 00.00 UTC.

498

CORAL REEFS: NATURAL DISTURBANCES

Figure 11.9. Map of Maumere Bay and the location of survey sites in November 1993. The survey applies only to coral communities at depths between 2-15 m. Numbers indicate severity of cyclone damage: 1) Low numerous branching and laminar Acropora s p p . , abundance of recruits; 2) Medium - visible damage to branching species, massive species present, numerous recruits; 3) High - branching species totally destroyed, massive species damaged but still in place, substrate of unconsolidated rubble and loose sediment is inhibiting recruitment; 4) Severe - most corals destroyed (95%-100%), very few recruits. Recruitment estimate is based on the abundance of juvenile colonies (2 uM or ug-at.l"), phosphorus acts as a crystal poison by inhibiting the formation of CaC0 3 crystals (Simkis 1964). The ecological effects of high nutrient concentrations are more difficult to interpret, since other environmental characteristics of the area (e.g., coastal currents, background levels, depth profile, etc.) must be considered. Nonetheless, sufficient data are available to conclude that an increase in nutrient concentrations above ambient will at some point lead to eutrophication. In this discussion we are interested in the effects of nutrient concentrations only slightly above the mean ambient levels, which are usually not investigated in field or labo-

546

CORAL REEFS: NATURAL DISTURBANCES

Figure 11.39. Large tabulate acroporids are a dominant component'of the 1988 lava flow coral community. This Acropora latistella is being used as a dinner plate by some reef inhabitants, most likely an octopus or a Napoleon wrasse (Cheilinus undulatus). Gunung Api, Banda Islands, May 1993 (maximum diameter 85 cm). Photo by Tomas and Anmarie

Tomascik.

ratory studies. For example, it has been demonstrated that under laboratory conditions, nitrogen and phosphorus have a measurable effect on the shell formation of Tridacna gigas (Belda et al. 1993). Working with juvenile giant clams, they were able to demonstrate that high nutrient levels (N and P) significantiy enhanced shell growth rates (i.e., skeletal extension), but at the same time also significantly reduced the weight of the shells. This implies that while the clams are growing rapidly under high nutrient concentrations, their shell densities are much lower. The implication is that lower shell densities may make the tridacnas more susceptible to predation by drilling gastropods, especially during their juvenile stage just after settling from the plankton. However, the nutrient (N and P) concentrations used in the study (i.e., 5-10 uM, N; and 2-10 uM, P) to elicit this response were well above the average ambient levels that the clams are likely to encounter on the reef. Unfortunately, there are no comparable studies using environmentally 'realistic' concentrations. What the study suggests, however, is that similar responses to elevated nutrient concentrations may be elicited in other calcifying photosymbiotic organisms, including the zooxanthellate reef-building corals. The above-average skeletal extension rates of scleractinian corals, as well as large shell size (maximum size 48 cm) of Tridacna squammosa (well above the average; R. Braley, pers. comm.) on the Banda Api lava flow, suggests that nutrients may be involved. During the Southeast Monsoon upwelling period, ambient sea surface P0 4 -P and NOs-N concentrations in the Banda Sea range between 0.11 to 0.56 uM and 0.30 to 3.52 uM, respectively (Institute of Marine Research 1973c; Wetsteyn et al. 1990), and are well above the average values in the Caribbean and elsewhere. During the Southeast Monsoon, all coral reefs in the Banda and Flores Seas are subjected to very similar nutrient conditions (Wetsteyn et al. 1990), with slighdy higher levels in the Arafura Sea. This difference is due mainly to the upwelling of 100-150-m-deep Arafura Basin water masses over the Arafura Shelf. According to Wetsteyn et al. (1990), this is an impressive phenomenon that brings nutrientrich water to benthic and pelagic communities hundreds of kilometres from the shelf break. Nonetheless, coral growth rates comparable to Banda Api corals have not been observed, even on some of the adjacent islands in the Banda Archipelago. This suggests that the effusion of hydrothermal fluids from numerous

GEOGRAPHIC DISTRIBUTION OF VOLCANOES

547

shallow-water hydrothermal vents around Banda Api may be a significant source of nitrogen and phosphorus to the reef community. For example, it has been demonstrated that rivers draining active volcanic regions in Costa Rica, Central America, carry high loads of dissolved nutrients, which significantly enhance the productivity of the receiving waters (Pringle et al. 1993). Therefore, it is suggested that hydrothermal fluids associated with geothermal processes of the Banda Api volcano may also be a significant source of nutrients. Unfortunately, data to demonstrate this probable connection are not available. However, thermal springs are a very common phenomenon on Java as well as other volcanic islands throughout the archipelago. Among the best known are the famous thermal springs of Cipanas in West Java, located at the foot of an old lava flow from Gunung Gede (Guntur). According to Hutchison (1981), Gunung Gede's lavas contain between 50%-61% Si0 2 and between 0.44%-0.88% KgO, whereas Banda Api's Si0 2 content ranges between 62.7%-67.8%, and K^O between 0.59%-0.89%; thus both are classified as andesitic volcanoes. While most lavas from Java volcanoes are quartz normative, the Banda Api lavas are entirely quartz normative (i.e., theoretically highly predictable mineral composition) (Hutchison 1981). It is, therefore, possible that the mineral and nutrient concentrations of the hydrothermal fluids associated with the geothermal activity of these two volcanoes may be similar. Much closer to Banda Api and situated on the Inner Banda Arc are three islands (i.e., Pantar, Serua and Damar) with active andesitic volcanoes. Hydrothermal activity along the coastline occurs on all three islands, and some of these hot springs were sampled during the Snellius-II Expedition. The expedition visited Banda Api in 1984, but samples from hydrothermal vents were not collected. However, the composition of the Banda Api hydrothermal fluids may be similar to those of Serua and Damar, since the hydrothermal vents are located near the beach (table 11.6). Unfortunately, nutrient analyses for N and P were not done during the Snellius-II Expedition, so possible nutrient budgets cannot be estimated. Therefore, hot spring data from Cipanas are included in table 11.6, since values for P and N can be approximated. The data clearly demonstrate that hydrothermal fluids are significantly enriched in nutrients and other mineral constituents; thus, it is likely that hydrothermal venting will have a measurable impact on the coral communities and the associated reef fauna and flora. Based on our observations thus far, this effect seems to be very beneficial. The seasonal monsoonal upwelling in the Banda Sea that occurs during the Southeast Monsoon (Wyrtki 1961; Nontji 1975; Birowo and Ilahude 1977; Boely et al. 1990; Ilahude et al. 1990; Wetsteyn et al. 1990; Zijlstra et al. 1990), and the continuous effusion of geothermal fluids from the volcano, are, however, not the only source of nutrients for the Banda Api lava flow coral community. As a result of submarine topography and tidal regime, the Banda Islands also experience daily tidally-induced upwelling events that may significantly influence the nutrient budget over the reef on a daily basis. Based on in situ temperature measurements taken over the reef slope during the upwelling events (with recorded temperature drops from 29.8°C to 26°C), and temperature and nutrient depth profiles taken during the Snellius-II Expedition (Wetsteyn et al. 1990), we have estimated that the upwelling water masses may originate from depths between 50 to 60 m. Nutrient concentrations measured at these depths in the Banda Sea range between 0.13 0.53 uM and 0.9 - 7.8 uM of P 0 4 and NO s , respectively (fig 11.40) (Wetsteyn et al.

548

CORAL REEFS: NATURAL

DISTURBANCES

1990). Note that these nutrient concentrations are not persistent, but occur in pulses. High ambient nitrate concentrations may, therefore, be one explanation for the recently demonstrated 8 N enrichment of Pontes spp. tissue samples collected directly from the Banda Api lava flow. These values are similar to the 815N values obtained from corals collected from a sewage-polluted area in the Maldives (Risk et al. 1993). It is worth mentioning that all Acropora colonies were very fragile and brittle, a condition exhibited by Tridacna gigas exposed to high N and N+P concentrations (Beldaetal. 1993). Coral colonization of lava flows has been previously reported only from Hawaii (Grigg and Maragos 1974), a region with much lower species richness when compared to the Banda Islands, as well as different environmental conditions. The scleractinian fauna of Hawaii is represented by only 12 scleractinian genera and 44 species (Grigg 1983; Dollar and Tribble 1993). In an important ecological study of scleractinian coral colonization rates of Hawaiian lava flows, Grigg and Maragos (1974) clearly demonstrated that Hawaiian scleractinian coral communities smothered by basaltic lava flows may take from 20 to over 50 years (i.e., depending on environmental conditions) to recover to their earlier species richness, percent cover, and diversity levels. The results of the Hawaiian study have been cited, and used frequently, to demonstrate the relatively slow recovery rates of coral communities following major environmental perturbations. These results fit well with the earlier observations made by Umbgrove (1947) at Krakatau. While the coral communities in Hawaii and Krakatau evolved under markedly different climatic and hydrological conditions, they nevertheless colonized a similar (i.e., chemical composition) substrate consisting predominantly of basaltic lava. Different envi-

Table 11.6. Major mineral and nutrient constituents of hot springs (hydrothermal vents) from Cipanas (West Java), and Pantar, Nila and Damar on the Inner Banda Arc. The Pantar and Serua hot springs are located on the beach; thus, thermal fluids are contaminated with seawater. Constituent range (mg.l 1 )

T(°C) pH

Cipanas

Pantar

Serua

Damar

44.0

80.0

70.0

81.0

2.2-2.8

6.5

6.5

6.5

Si0 2

132-157

162

197

242

S0 4

255-407

1200

1652

270

0

0

488 501 278 5612

53 12 42 114

NH3

HP0 4 P Al Fe (total) Fe (inorganic)

Ca Mg K Na

5-11

0.28-3.00 0.43 - 0.97 22.20 - 65.59 4-20 4-18

65-111 17-30 19-29 37.61-76.95

Sources: van Bemmelen 1949; Poorter et al. 1989.

0.7

330 550 190 3950

GEOGRAPHIC DISTRIBUTION OF VOLCANOES

549

Figure 11.40. Vertical depth profiles for temperature (T°C), nitrate (N0 3 -N uM) and inorganic phosphate (P0 4 -P uM), Banda Sea. The stations were in the vicinity of the Banda Islands during the non-upwelling season in April (i.e., Northwest Monsoon), and the upwelling season in September (i.e., Southeast Monsoon). Non-upwelling: April 8, 1970 (Station 16; Cruise No. 1). Upwelling: September 22, 1972 (Station 25; Cruise No. 10). Source: Institute of Marine Research 1971a, 1973b.

ronmental conditions that occur in Sunda Strait (i.e., location of Krakatau) and the Banda Sea (i.e., location of Gunung Api), to a great extent explain current differences in coral community structure and reef development. However, it is suggested that substrate type has been the major environmental factor that has considerably slowed down the reef development at Krakatau. Unlike Gunung Api, where solid andesitic lava is the primary substrate, loose basaltic lava rock and thick tephra deposits that form high eroding cliffs clearly predominate at Krakatau. Thus, coral communities at Krakatau have little chance to become established, a very similar scenario that also exists along the south coast of Gunung Api. The Gunung Api eruption in 1988 created a natural experiment where the colonization of an andesitic lava flow could be monitored for the first time. A 1993 study of the lava flow (five years after the eruption) provided first quantitative evidence that, under natural conditions (i.e., unaffected by anthropogenic activities) , coral communities have a remarkable healing capacity to recover from major environmental perturbations (Tomascik et al. 1996). Within five years of the eruption, the new coral community (the embryonic reef) that became established on the lava flow has surpassed the earlier and adjacent coral communities, which were not affected by the eruption, in terms of species richness, abundance, and cover. The rapid colonization and succession may be related to the above-average coral growth rates, and to the apparent preference for andesitic lavas by settling coral planulae. The contrasting results of the Hawaiian and Gunung Api studies are mainly a result of different environmental conditions, frequency of local disturbances (Dollar and Tribble 1993), lava type, biogeographic setting, the diversity of the local

550

CORAL REEFS: NATURAL DISTURBANCES

and regional species pool, and perhaps anthropogenic influences. The common feature of most of the young Indonesian volcanoes is that their lavas are mostly andesitic. This seems to be the common denominator in all instances where we have observed vigorous coral reef development to be occurring at the present. Table 11.7 lists the major elements of lavas from all active volcanoes where extensive reef development was recorded. The data presented in table 11.7 clearly demonstrate that major elemental differences exist between the Banda Api and the Hawaiian (Kilauea) lavas. The 1993 Banda Api study provides sufficient grounds to seriously question the generality of the widely held view that coral reefs do not develop on active volcanoes (Davis 1928). While it is true that Banda Api does not have a continuous fringing reef, long stretches of the coastline, especially along the northeast coastline, are fringed by a well-developed fringing reef. The absence of a well-developed fringing reef along the south and southwest coast of the volcano is primarily a result of unstable substrate, which consists mainly of loose pyroclastics and large lava boulders that are severely scoured by loose volcanic ash particles. However, Banda Api is an integral part of a three-island complex (i.e., Banda Naira, Banda Besar and Gunung Api - not including the smaller islands), which is in fact almost totally surrounded by a coral reef.

Table 11.7. Elemental composition (wt%) of lavas from the Sangihe and the Inner Banda Arc volcanoes that are known to support vigorous reef development. Location/Name

Si0 2

Ti0 2

AI 2 0 3

FeO

MgO

CaO

Na 2 0

K20

P205

Sangihe Arc Awu (Sangihe) Kahakitang Siau Makelehi Ruang ManadoTua

54.14 51.14 53.58 59.16 52.37 58.15

0.80 0.78 0.79 0.53 0.69 0.58

19.29 17.58 18.67 20.00 20.05 18.87

7.99 8.31 8.74 5.05 8.36 6.26

4.10 7.42 4.49 2.01 4.92 2.91

9.44 10.87 9.34 6.90 10.05 7.60

3.19 2.71 3.02 4.19 2.71 3.24

1.16 0.96 1.07 2.14 0.54 1.44

0.14 0.14 0.16 0.23 0.11 0.23

Average

54.76

0.70

19.08

7.45

4.31

9.03

3.18

1.22

0.17

Banda Arc Banda Api Banda Naira Banda Besar Manuk Serua Nila Teun Damar

64.44 55.26 55.10 57.49 56.45 57.08 57.59 56.16

0.98 1.23 1.31 0.76 0.69 0.59 0.64 0.66

14.56 16.65 16.35 17.50 17.89 17.29 18.38 17.78

7.27 10.44 11.65 7.39 7.71 7.91 7.03 8.06

1.76 3.50 3.19 4.01 4.58 3.80 3.23 3.82

4.91 8.34 8.04 8.01 8.76 8.50 7.50 8.11

4.10 3.52 3.36 3.47 2.50 2.53 3.19 2.70

0.89 0.65 0.55 1.07 1.14 2.01 2.10 2.37

0.25 0.17 0.18 0.13 0.10 0.10 0.17 0.15

Average

57.45

0.86

17.05

8.43

3.49

7.77

3.17

1.35

0.16

Total average

56.29

0.79

17.92

8.01

3.84

8.31

3.17

1.29

0.16

Hawaii*

50.23

2.57

13.42

12.06

7.28

10.98

2.31

0.50

0.26

* Source: Garcia et al. 1992. Sources: van Bergen et al. 1989; Jezek et al. 1981.

GEOGRAPHIC DISTRIBUTION OF VOLCANOES

551 Figure 11.41. During volcanic eruptions burial of fringing reefs may occur in coastal areas downwind from the eruption. Large sponges such as this Petrosia testudinaria may recover from partial burial. Southwest coast of Ruang Island, Sangihe Archipelago. Photo by Tomas and Anmarie Tomascik.

High-diversity coral communities, such as those found on the Banda Api lava flow, are not unique to the "Spice Islands". Similar high-diversity and high-cover coral reef communities that now form well-developed fringing reefs have been observed around other active volcanoes in the Banda and Sulawesi Seas (e.g., Ruang Island - with the last eruption in 1949). These observations clearly demonstrate that volcanoes do not have to be extinct or dormant before coral communities, and thus a fringing reef, can develop. Here again the growth of the reefs was more vigorous than I had expected to find on the slope of an active volcano. . . Although Gn. Api [north of Wetar] cannot therefore be cited as a proof of the theory that reefs can grow on an active or only recently extinct volcano, it certainly agrees with this hypothesis.—KUENEN 1933

Kuenen's earlier, and often overlooked, suggestion has found strong support in our quantitative study of Banda Api as well as from our qualitative observations on many other volcanic islands with active volcanoes. Thus, contrary to the long-held view, the Gunung Api study in the Banda Islands suggests that intermittent eruptions may actually promote local coral diversity, and that hydrothermal activity associated with geothermal volcanic processes may promote, in an unknown fashion, coral growth rates (Tomascik et al. 1996). Nonetheless, Pelean eruptions without significant lava flows produce considerable amount of pyroclastic "fallout" (i.e., volcanic ash, tuffs, etc.) that may smother considerable sections of coastline, including fringing reef communities (fig.. 11.41). Clearly, this area offers exciting new research opportunities that can provide us with much-needed answers to some very complex questions. Banua Wuhu (Mahengetang) Underwater Volcano Hydrothermal venting that was observed in Banda Api is not a unique phenomenon in the archipelago, and most likely occurs widely wherever there are active volcanoes. The Galapagos Islands are famous, at least in sport-diving circles, for their gasbubbling, hot-water vents associated with the volcanic nature of the islands. The

552

CORAL REEFS: NATURAL DISTURBANCES

Box 11.4. Corals from the Banda Sea as environmental recorders. J. Heikoop and M. Risk, Department of Geology, McMaster University, Hamilton, Ontario, Canada. During recent years, the flourishing reefs of the Banda Islands, Indonesia, have been affected by the 1988 eruption of Banda Api and have experienced oceanographic variability associated with their presence within the Western Pacific Warm Pool. Massive corals collected from reefs offshore of the Banda Islands contain physical and chemical records of these phenomena in their aragonite skeletons. Banda Api is a small volcanic island forming part of the Banda Arc, a collision zone between the Eurasian and Indo-Australian tectonic plates. The volcano last erupted May 9, 1988, depositing up to 1 m of ash on the surrounding coastline. Local coral mortality resulting from the eruption was extensive, particularly among branching species. In May, 1992, four heads of the coral Porites lobata were collected just offshore of Banda Api to determine if a record of this eruption might be contained within their skeletons. It was thought that ash from the eruption may have been incorporated into the skeletons and that this material would be revealed by petrographic investigation. When the coral heads were cleaved open, two interesting features were immediately evident. All four corals contained extensively bioeroded death/regrowth surfaces through which volcanic ash had infiltrated into the skeletons (fig. 11.42). At the same level within the skeleton as the death surfaces, a thin, black, iron-rich precipitate we have termed the "Banda Band" was also found (fig. 11.42). This band quickly oxidized to an orange colour upon exposure to the atmosphere. Using X-radiography and stable isotope chronologies, the contemporaneous Banda Band and death surfaces in each coral were found to have formed at the time of the May 1988 eruption of Banda Api (Heikoop et al., in press). Death surfaces formed over parts of the skeleton where ash from the eruption could not be quickly cleared away by wave energy or by coral sediment rejection. These surfaces were subsequently bioeroded and infiltrated by small ash particles prior to tissue expansion and regrowth over damaged areas of the coral colony. The origin of the Banda Band is thought to be related to local hydrothermal activity which has been ongoing near the collection site since the eruption. A large area of hydrothermal venting deposits iron oxides on surrounding rocks and sediments close to the collection site. It is likely that enhanced activity, coinciding with the 1988 eruption, was responsible for deposition of the chemical precipitate onto the coral skeletons. It is unlikely that any discrete record of the 1988 eruption will be preserved within surrounding sediments, due to the relatively moderate size of the event and the occurrence of extensive biological and hydraulic sediment reworking. It is quite possible, however, that century-old massive heads in this or other similarly active volcanic/hydrothermal areas may record one or more of these geologically instantaneous events and that similar records could be found in fossil corals. The hydrothermal venting around Banda Api has also been of interest as a source of possible nutrient input to the reef system. This work has focused on a number of coral tissue samples collected from small colonies of Porites lobata growing directly on the Banda Api lava flow. These corals had quickly colonized the lava flow which had entered the sea on the north coast of the island during the 1988 eruption (see Banda case study this section). Hydrothermal venting occurs adjacent to this lava flow. The coral tissues were analyzed at McMaster University and were found to be enriched in the heavy isotope of nitrogen. They were, in fact, among the heaviest tissues (815N= 6.8 ±0.5 %0; n=9) from among the more than 200 samples we have measured from reefs around the world. Carbon isotopic values were typical of other shallow-water corals we have measured.

GEOGRAPHIC DISTRIBUTION OF VOLCANOES

553

Box 1 1 A (Continued.) Figure 11.42. Photo of a cleaved coral head collected offshore of Banda Api. Large arrow indicates "Banda B a n d " . Small arrow points to bioeroded death surface through which ash (a) has infiltrated into the skeleton. Scale 1:2.

Figure 11.43. 5180 record reconstructed from a coral colony collected off the west coast of Run Island in the Banda Sea. Small drilled samples were taken sequentially down the major growth axis of the coral and analyzed on a VG SIRA mass spectrometer to obtain this record. Timing of the record is indicated at the top of the graph. Chronology was established by X-radiography.

554

CORAL REEFS: NATURAL DISTURBANCES

Box 11.4. (Continued.) Heavy nitrogen isotopic signatures in corals may reflect high light levels (Muscatine and Kaplan 1994; Heikoop et al., in prep.), high nutrient concentrations or an isotopically enriched source of dissolved inorganic nitrogen (DIN) (Risk et al. 1993). All of these factors may be important in this particular setting. Higher concentrations of DIN or an isotopically enriched nitrogen source could be provided by local hydrothermal venting. It is also possible that these corals may be affected by the endo-upwelling mechanism proposed by Rougerie and Wauthy (1993). An active volcanic system such as Banda Api could be expected to create a geothermal gradient, causing upward convection of deep oceanic waters through the volcanic substrate upon which the reef is constructed. Reefs developed around the Banda Islands are not only influenced by volcanism. These reefs are also flourishing in the warmest body of water on Earth, the Western Pacific Warm Pool, with temperatures generally greater than 28°C (Yan et al. 1992). During El Nino/Southern Oscillation (ENSO) events the warm pool migrates eastward to be replaced by somewhat cooler waters. ENSO events are also associated with severe drought in Indonesia. Given the severe impact of ENSO on eastern Pacific sea surface temperatures and meteorological conditions around the globe, proxy records of this pan-Pacific phenomenon are needed for climate reconstruction modeling. Records from the warm pool are of particular importance in this regard, as sea surface temperature and salinity anomalies in this body of water are potentially responsible for causing the westerly wind bursts which signal the onset of an ENSO event (Enfield 1989). An oxygen isotope record from a Pontes lobata skeleton is presented in figure 11.43. This preliminary work, from a young coral colony collected at Run Island, shows several features which may be of importance. The trend towards slightly lower 8180 between 1984 and 1990 could be due at least in part to a warming trend in the warm pool (Heikoop and Risk 1994). Evidence for warming is contained in both satellite data (Yan et al. 1992) and ship-collected COADS data for this time period. The isotopic trend is interrupted in 1991-1992, which was the beginning of a large prolonged ENSO event. The slightly higher 5180 values associated with this event could be due to a combination of lower SST and increased 5180 of seawater associated with drought conditions (Heikoop and Risk 1994). However, Yan et al. (1992) point out that significant cooling may have occurred as a result of the 1991 eruption of Pinatubo in the Philippines, which released large quantities of aerosols into the stratosphere, thus reducing insolation. There is only very weak evidence for the presence of the moderate 1986-1987 ENSO event in this record. The isotopic record does not extend back far enough in time to record the severe 1982/1983 ENSO event. Isotopic reconstructions from longer coral records have the potential to greatly increase understanding of ENSO in this area. Corals are proving to be very important proxy recorders of both volcanism and oceanographic variability in the marine realm of the Banda Islands. With proper calibration, such records could be of great help in filling in gaps in our knowledge of past natural phenomena associated with this region.

GEOGRAPHIC DISTRIBUTION OF VOLCANOES

555

same phenomenon, but on a much grander scale, was observed on the Sangihe Ridge, which stretches from North Sulawesi to Mindanao in the Philippines. The ridge is a volcanic arc delineating the western border of the Molucca Sea, and is part of a unique arc-to-arc collision complex (Jezek et al. 1981). The volcanic arc, about 550 km long and up to 70 km wide, has a series of volcanoes, spaced 30-90 km apart, that lie between 130-180 km above the centre of a west-dipping Benioff zone that is about 650-km-deep (Jezek et al. 1981). More than 20 of the islands, along the southern and central section of the volcanic arc, are young Quaternary volcanoes. All of these have well-developed fringing reefs that in most instances encircle the entire island, with flourishing coral communities characterized by high diversity and cover. In all, there are 13 active volcanoes along the ridge, including an active submarine volcano. The southernmost volcano on the Sangihe volcanic arc is Unauna, in the Gulf of Tomini, which last erupted in 1983. Most of the islands in the northern region of the arc (i.e., north of Sangihe Island) are capped by thick coral reef deposits, in addition to being currently surrounded by extensive fringing reefs (e.g., Kawio Islands). The Sangihe Ridge has a high diversity of coastal and oceanic habitats, such as, for example, Pasige Island, which is an oceanic platform reef (almost an atoll - but not quite) with an extensive mangrove forest in one part of the shallow (50 cm deep) deposit of ferriferous ooids/pisoids on the west coast of Mahengetang Island. Note the absence of calcareous material and the homogeneous nature of the ferriferous deposit. B) The extensive ooid/pisoid fields are major hydrothermal and gas venting areas. Photos by Tomas and Anmarie

Tomascik.

EARTHQUAKES Earthquakes are a common occurrence in many parts of the world. Examination of the main active earthquake belts reveals that, almost without exception, major earthquake activity occurs along the margins of interacting tectonic plates, while they rarely, if ever, occur in many other regions. On a geological time scale the Indonesian Archipelago is relatively young, formed by the interaction of three major tectonic plates, the Australian-Indian Plate to the south, the Pacific Plate to the east and the Eurasian Plate to the north, as well as a number of smaller plates (see chapter 2). Major interactions occur along all active margins, for example, along the Sunda Arc where the Indian-Australian Plate is being subducted under the Sunda Shield portion of the Eurasian Plate. Thus, Indonesia is one of the most earthquake-prone regions in the world. Fortunately, the most destructive earthquakes (i.e., shallow earthquakes) have their epicenters located in offshore oceanic regions, or in relatively underpopulated areas (Katili 1985), otherwise the damage sustained would be much greater. However, with the predominant offshore location of major earthquake epicenters, the probability of a tsunami impacting upon the shorelines increases as well. According to Katili (1985), 10% of the world's earthquake activity occurs in the Indonesian Archipelago, whose total area of 5.8 x 106 km 2 (including the Exclusive Economic Zone - EEZ) (Uktolseya 1990) represents roughly 1.1 % of the earth's surface. Table 11.8 provides a summary of the Indonesian seismo-tectonic zones. As was clearly pointed out in the RePPProT (1990) report, nearly 90% of the Indonesian population lives in regions where earthquakes with a magnitude of at least 8.0 on the Richter scale can be expected, and Jakarta is within one of these danger zones. The disastrous consequences of the earthquakes in Managua,

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Nicaragua, and Mexico City, Mexico, two cities built in areas with not-too-dissimilar geologic histories, should not be forgotten. Recent tragedies in Flores and South Sumatra serve to remind us of the dynamic nature of the archipelago. Shallow-water earthquakes (earthquake foci 50 cm.sec"1) and frequent upwelling (i.e., tidally induced) along the entire seaward margin of the Berau Barrier Reef are the two key factors most likely responsible for the development, and the vigorous nature, of the reef complex in what seems to be a major deltaic depositional environment. Reef development in this region, what Netherwood and Wight (1993) described as a delta-front setting, dates as far back as late Miocene to early Pliocene. The foundation of most shelf reefs consists of various continental crust sedimentary rocks and sandy deposits, that are part of the normal series of erosional products originating from the adjacent continental blocks forming the continental

SHELF REEFS

585 Figure 12.4. The Berau Barrier Reef system off the coast of East Kalimantan. Note the depth of the narrow lagoon and numerous patch reefs. The mainland is fringed by a wide fringing reef dominated by a mixed community consisting of massive Pontes spp., Goniastrea spp. and soft corals (e.g., Alcyonium spp., Lobophytum spp., Sinularia spp., Sarcophyton spp.). The net flow through the channels throughout the year is in a north-to-south direction. Tidal currents have an important local effect on reef zonation and development. Intertidal reef flats are lightly stippled, mainland Borneo is heavily stippled. Drawing by B. Rahmad.

slope (Fagerstrom 1987; Longhurst and Pauly 1987). Shelf reefs, with few exceptions, are not areally limited and are usually more numerous and attain greater dimensions than oceanic reefs. For example, the Pulau Panjang Reefjust north of the Berau River delta is over 30 km long with an area of about 150 km2. About 40 km to the south of Berau delta lies a vast patch reef, Karang Besar (Big Reef), with an estimated area of about 300 km2. Whether this massive reef complex has a lagoon is not known. There is considerable variation in reef types along the inshoreto-offshore axis. While in some regions coral reefs are present almost continually from the shoreline (i.e., fringing reefs) to the seaward margin of the shelf (e.g., Watampone Shelf, Spermonde Shelf), in others regions, the reefs form a welldeveloped barrier along the seaward margin (i.e., 200 m isobath) of the shelf, but they are distributed only intermittentiy in the inshore regions (e.g., North Sumatra, Gulf of Tomini). In general, coral reefs along the seaward margins of shelves are different from the inner shelf reefs. For example, the outer shelf reefs of the Spermonde Barrier Reef or the Great Sunda Barrier Reef are much bigger, longer and, in general, more vigorous. The Balabalagan Reef located at the north margin of the Great Sunda Barrier Reef is located about 120-150 km offshore, and runs almost uninterrupted for 75 km along a southeast-to-northwest axis at the seaward margin of the Sunda Shelf. However, nothing is known about its ecology or geology. The vast majority of seaward shelf reefs in the archipelago are located in upwelling regions. One of the major benefits of upwelling to coral reefs is that the nutrientrich waters may stimulate organic production of CaCO s . It is important to stress the

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fact that there are in effect two types of upwellings in the Indonesian Archipelago. The classical upwelling in both the Flores and Banda Seas, as described by Nontji (1975); Boley et al. (1990); Ilahude et al. (1990); Wetsteyn et al. (1990); and Zijlstra et al. (1990), is generated by the monsoonal winds during the Southeast Monsoon when strong southeasterly winds drive the surface water masses in a westerly direction. The second type of upwelling which we observed throughout the archipelago is tidally induced. These events have a significant local effect and may be described as high-frequency, low-duration phenomena. Fagerstrom (1988) suggested that cooler seawater temperatures associated with upwelling may stimulate inorganic production of CaCO s through loss of dissolved C0 2 . Whether this process is of any significance in reef development in upwelling regions remains to be determined. It is generally assumed, especially in Indonesia, that coral reefs on the Sunda Shelf are depauperate and not well developed. However, Fagerstrom (1987) and Guilcher (1988) point out that shelf reefs, in general, exhibit greater diversity (geomorphological and bio tic) than oceanic reefs, mainly because of a greater complexity of factors that control their development. With the exception of extensive research conducted on the offshore patch reefs of Kepulauan Seribu, a shallowwater back-arc environment, not much information is available from other regions of the Sunda Shelf. Assessments of fringing reefs in Sunda Strait, Riau Islands, Natuna Islands, Belitung Islands, Karimunjawa Islands, Jepara and Kangean Islands indicate that coral reef systems in these areas may have been heavily damaged by anthropogenic activities, thus little can be said about their original diversity. Hopefully, new research efforts along the west coast of Sumatra and the Mentawai Islands, will provide new and much-needed information on the continental shelf reef systems facing the Indian Ocean (A. Kunzmann, pers. comm.). The most in-depth study of shelf reef systems, other than Kepulauan Seribu, was conducted by Moll (1983) who studied the zonation and diversity of scleractinian coral communities on the reefs off the southwest coast of Sulawesi (i.e., Spermonde or Sangerang Shelf). In a parallel geomorphological study of the reefs on the east-west tilting Spermonde Shelf, De Klerk (1983) described a variety of reef types (e.g., crescent-shaped, oblong, circular, etc.) and suggested that they were of Holocene age. In an earlier study of the Spermonde reefs, Umbgrove (1930) discussed the role of monsoonal winds and their role in the development of patch reef systems and coral cays. The significant role of the Northwest Monsoon in the geomorphology of coral cays in the Spermonde Archipelago is clearly evident from well-developed shingle ramparts along the western margins of the reef flats (de Klerk 1983). The first detailed taxonomic account of coral communities on the Spermonde Shelf was that of Wijsman-Best (1977), which was later expanded upon by Moll (1983), who described a total of 262 scleractinian species belonging to 78 (sub-) genera.

CLASSIFICATION OF S H E L F R E E F S The development and geomorphology of shelf reefs is a science in its own right, and well beyond the scope of this book (for review see Hopley 1982, 1983). While the Indonesian Archipelago has the largest shelf area in the world, detailed geo-

CLASSIFICATION OF SHELF REEFS

587

morphological studies of its reefs have been limited to the relatively small Seribu Platform on the Sunda Shelf (Umbgrove 1928, 1929b, 1939, 1947; Verwey 1931b; Hardenberg 1939; Verstappen 1954; Scrutton 1976a,b; Burbury 1977; Cook and Saito 1984; J o r d a n 1985; Ongkosongo and Sukarno 1986; Stoddart 1986; Ongkosongo 1988; Brown 1991; Park et al. 1992; Jordan et al. 1993; Longman 1993). With a few exceptions, the emphasis has been either on the formation of coral islands (i.e., coral cays), which are associated with many shelf patch reef systems, or on reef structure and sedimentology (Park et al. 1992; Jordan et al. 1993). Genetic classification of Indonesian shelf and oceanic reefs has been restricted to the classical Darwinian scheme based on the early works of Molengraaff (1922, 1929), Kuenen (1933a, 1947) and Umbgrove (1947). However, no attempt has been made to develop a classification system for the great diversity of reef types found within the classical Darwinian fringing, barrier and atoll reefs. The plethora of ambiguous and redundant terms that has pervaded Indonesian coral reef literature is a reflection of this serious lack of information, and is an issue that needs to be addressed in a systematic manner. The most recent comprehensive treatment of this complex subject is by Hopley (1982,1983), who proposed a general classification system for shelf reefs of the Great Barrier Reef, which, however, has global applicability. Hopley's (1982, 1983) classification scheme expanded upon earlier works of Fairbridge (1950, 1967) and Maxwell (1968, 1970), who based their schemes on the premise that the reefs of the Great Barrier Reef were of Holocene origin, and were a function of organic and sedimentary growth in response to prevailing wind and wave conditions under eustatic stillstand. Hopley (1982, 1983) included in his classification three additional factors, namely, the depth and morphology (e.g., size and lagoonal characteristics) of pre-Holocene antecedent platforms, the nature of the sea-level rise during the Holocene, and the rate of production of calcium carbonate. Because of the complexity and extent of the problem (Stoddart 1978), as well as the lack of sufficient information on recent Indonesian shelf reef systems, a meaningful genetic classification scheme for Indonesian shelf and oceanic reef types is not currendy possible. The objective of this cursory review is to streamline some of the key terminology used in Indonesian scientific literature, and caution on the indiscriminate use of ambiguous terms from the past. A good starting point is Stoddart's (1978) review of descriptive reef terminology. The ambiguous usage of the term atoll and its many derivatives needs to be addressed. From a Darwinian viewpoint, atolls are reefs linked to fringing and barrier reefs through an evolutionary sequence driven by the subsidence of their foundations, which may be volcanic or sedimentary in origin. However, the atoll is frequently used to describe any annular reef with a central lagoon (colour plate 12.1). The term is especially prevalent in popular coffee-table literature, where the Kepulauan Seribu lagoonal reefs are often referred to as atolls. Stoddart (1978) provides a detail account of the various terms that have been applied to grossly similar forms. For example, the term pseudo-atoll was used to describe annular reefs not formed by subsidence. Another term that is frequently used in Indonesia to describe an annular patch reef with a well-developed moat or a shallow lagoon is miniatoll (Longman et al. 1993). Scoffin and Stoddart (1978) discussed the development of a miniatoll in the context of reef types found within lagoons of large shelf reefs. Hopley (1982) also used the term to describe a type of patch reef found in the lagoons of shelf reefs on the Great Barrier Reef. A miniatoll is a

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Figure 12.5. A visual comparison of form between a mature lagoonal shelf reef and an atoll. A) Pulau Ayer is a vegetated cay on the leeward side of a medium-sized (2.5 km x 1 km) mature lagoonal shelf reef in Kepulauan Seribu (looking towards the east). The lagoon is less than 8 m deep. The reef is exposed along its southern margin (right), indicated by a white breaker line. B) Karang Mariane is an unusual atoll in the Flores Sea, rising from a depth of about 600 m (looking southwest). Karang Mariane may be a double atoll. Note the narrow (c. 500 m wide) submerged southern rim (far left) that is 2-4 m below sea level. The southern rim and the north atoll (foreground) enclose a lagoon 50 m in depth. The depth of the lagoon on the northern atoll is not known and may be in a stage of infilling. A. Photo courtesy of Ft. K. Park, Maxus southeast Sumatra, Jakarta; B. Photo by Tomas and Anmarie

Tomascik.

patch reef where the central area is depressed, by rapid bioerosion, relative to the rim (Scoffin and Stoddart 1978). Since the term miniatoll is already in use for a fundamentally different reef type, it is, therefore, suggested that the term miniatoll should not be used to describe the lagoonal patch reefs in the southern section of the Kepulauan Seribu. Figure 12.5 illustrates the gross geomorphological similarity of a mediumsized lagoonal reef, Pulau Ayer, in the Kepulauan Seribu Complex and an atoll (Karang Mariane) in the Flores Sea. Pulau Ayer (fig. 12.5A) (also known as Pulau Air; Chart # 414-KK) is a vegetated coral cay that developed on the leeward side of a medium-sized lagoonal reef. While the reef may resemble an atoll in general appearance, it originates from an upward growth of coral from a tectonically stable antecedent platform that was flooded during the Holocene transgression. Park et al. (1992) have shown that reefs on the Seribu Platform grew most of their bulk between 8000 and 4500 yrs B.P. with the rising sea levels. Thus the reef is a result of an eustatic change, and largely unrelated to diastrophism. However, Jordan et al. (1993) suggested that differential subsidence rates between the north and south sections of the Seribu Platform account for the differences in reef geomorphology. According to Jordan et al. (1993), the lagoonless, steep-sloped reefs at the north section of the island chain appear to be growing upwards, almost vertically, as if trying to keep pace with sea-level rise. In contrast, reefs at the south part of the Seribu Platform have reached sea level and are now expanding laterally. The presence of lagoons is explained by their inability to keep pace with higher rates of

FRINGING REEFS

589

sedimentation associated with reef fades at the reef margin. Hopley (1982) suggested that the morphology of the pre-Holocene antecedent platform, combined with concentration of growth around its periphery, are largely accountable for the formation of lagoonal reefs. The formation of atoll lagoons is still a hotly debated topic, nonetheless, they are usually much deeper; for example, the lagoon of Mariane Atoll (fig. 12.5B) is in excess of 40 m. Before a meaningful classification scheme can be developed for the Indonesian reef systems, a great deal of geomorphological work needs to be conducted. In the meantime, the classification system developed by Hopley (1983) would seem to be the most appropriate. For example, based on the gross geomorphology of the reef and surrounding bathymetry as well as the history of Holocene reef development in Kepulauan Seribu, Pulau Ayer Reef (fig. 12.5A) can be classified as a mature lagoonal continental-shelf reef with a vegetated cay. Hopley's (1983) classification scheme for shelf reefs is presented in figure 12.6.

FRINGING R E E F S One would expect that the term fringing reef (Darwin 1842) is self-explanatory, referring to coral reefs skirting continental and island (oceanic and shelf) shorelines. However, as Stoddart (1978) points out: " . . .it is frequently discovered that similar forms can have dissimilar origins, and that form alone is an ambiguous guide to genesis". This observation seems to apply to fringing reefs, frequently considered as the simplest of reef types. The reefs illustrated in figure 12.7 show similar forms, yet their geneses are fundamentally different. At a glance, both seem to illustrate the classic example of fringing reefs skirting a shoreline. The reef shown in figure 12.7A is a classical (i.e., in a Darwinian sense) fringing reef that evolved because the shelf of the island provided a platform that was suitable for colonization by a variety of reef-building organisms (i.e., scleractinians and coralline algae). While the reef is a product of abiotic and biotic factors that have acted in synergy throughout the reefs' development history, the reef is fundamentally a product of the island shelf. In very simple terms, no island, no reef. In contrast, the reefs skirting the two islands in figure 12.7B cannot be considered as fringing reefs, since the two islands are in fact vegetated coral cays that are products of the reefs themselves. In this case, the reefs are part of an outer rim of Taka Bone Rate Atoll in the Flores Sea. The coral cays are a product of reef growth, environmental conditions (e.g., wind, waves, currents, etc.) and sea level history. Other reefs that have frequently been misinterpreted as fringing reefs are the patch reefs of Kepulauan Seribu, many of which have well-developed coral cays (fig. 12.8). Siswandono et al. (1993) have recently used LANDSAT MSS images to map a number of coral reef areas in Lombok, West Nusa Tenggara. They classified the extensive reef systems skirting Gili Air, Gili Meno and Gili Trawangan as fringing reefs. During our visit to Gili Air and Meno we found no evidence of any extrusive sedimentary or volcanic sediments which would suggest that the initial reef development occurred around pre-existing islands. The bathymetry of the shelf and the absence of non-coralline sediments on the three islands suggest that they may be undergone considerable tectonic uplift since their formation. If this is correct, then the reef classification

590

CORAL REEFS: G R O W T H AND DEVELOPMENT

Figure 12.6. Schematic outline of a genetic classification of continental shelf reefs. The classification is given for three size classes of antecedent platforms since the flooding by the Holocene transgression. 1 (A-F) Type V reefs: Large-sized antecedent platform (> 3.25 km width). 2 (A-G) Type 2' reefs: Medium-sized antecedent platform (1.75-3.25 km width). 3 (A-F) Type 3' reefs: Small-sized antecedent platform (< 1.75 km width). From Hopley 1982, 1983.

FRINGING REEFS

Figure 12.6. (Continued.)

591

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CORAL REEFS: G R O W T H AND DEVELOPMENT

Figure 12.6. (Continued.)

FRINGING REEFS

593

Figure 12.7. A comparison of similar yet fundamentally different reef types that are frequently lumped together. A) An example of a classical Darwinian fringing reef along the west and north coast of Pulau Pasi, Flores Sea. Pulau Selayar is in the background (looking northeast). The fringing reef has a well-developed boat channel (sensu Guilcher 1988) that is 1-2 m deep and supports extensive coral growth. B) The Pasitelu Island group on the southwest rim of Taka Bonarate Atoll. Pulau Tengah to the west (left of centre) and Pulau Raja to the east (upper right) are vegetated coral cays (c. 1 km x 0.4 km) at the centre of two rim reefs that form the southwest margin of Taka Bonarate Atoll. Photos by Tomas and Anmarie

Tomascik.

Figure 12.8. A typical island in the northwest sector of Kepulauan Seribu with a skirting reef. This type of reef is often misclassified as a fringing reef. The reef is a small platform reef capped by a vegetated coral cay. Compared to the coral cays further south, coral cays in the north sector of the island chain cover a much greater portion of the reef. Photo courtesy of R. K. Park and Maxus southeast Sumatra, Jakarta.

becomes confusing, since while the reefs fringing the three islands are similar in form to the fringing reefs along the mainland (e.g., Tanjung Sirrah), they are nonetheless genetically distinct. Without information on the internal structure of the islands and reefs, as well as the chronology of sea level change, classification of these reef types is purely speculative.

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OCEANIC REEFS Based on extensive geomorphological research in the Pacific Ocean, a number of different types of oceanic reefs are recognized (Scott and Rotondo 1983). The evolution of reefs on the Pacific Plate is geologically distinct from the evolution of reefs in the Indonesian Archipelago. Whereas the origins of the majority of reefs in the archipelago are related, in one way or another, to tectonic collisions between different tectonic plates, the evolution of Pacific Plate reefs results mainly from the slow motion of the basaltic Pacific Plate, in a northwest direction from its origin at the East Pacific Rise. The East Pacific Rise is a major sea-floor spreading centre, a type of conveyor belt, where upwelling magma from the asthenosphere is continually forming new lithosphere as it moves away from the spreading centre. At the other end of the conveyor belt are the Kermadec and Tonga Trenches (i.e., subduction zone) in the central south Pacific, and the Mariana, Japan, Kuril and Aleutian Trenches in the northwestern and north Pacific. The East Pacific Rise is a continuous formation about 2 to 4 km wide with an average depth of 2.7 km. As the Pacific Plate moves along in a northwest direction, it passes over a number of "hot spots", which are theoretical centres of volcanic activity beneath the lithosphere (i.e., cmstal plate). For example, the Hawaiian island chain is a result of the movement of the Pacific Plate over such a hot spot, which can be considered as a magmatic anomaly (Grigg 1982; Scott and Rotondo 1983). The Indonesia Archipelago, while tectonically active, has no such counterpart. Figure 12.9 presents a simplified version of the evolution of major reef types on the Pacific Plate. The Indonesian Archipelago, nonetheless, contains a variety of reef types that are structurally similar to Pacific Plate reefs, even though the geologic processes involved in their evolution may be different. Review of Indonesian coral reef literature and specific site visits to a number of geologically distinct regions of the archipelago revealed that scleractinian coral communities on oceanic reefs flourish in a variety of environments. One of the main characteristic features of oceanic reefs is that they are located away from major influences of continental processes (e.g., land runoff, siltation, turbidity, etc.). This is a characteristic environmental setting of many atolls in the Pacific, which are located thousands of kilometres from continental landmasses. This feature, however, is somewhat blurred with regards to the oceanic reefs in the Indonesian Archipelago, since many so-called oceanic reefs that rise from great depths are nonetheless located in close proximity to large islands. The geomorphology, and the ecosystem structure, of oceanic reefs is largely determined by the shape, size and physico-chemical composition of the foundation upon which they have become established. Simply put, the larger the original foundation, the larger the reef capping it. Taka Bonarate, which is the largest atoll in the archipelago, is a good example of this generalization, since it developed on top of an extensive ridge system (Kuenen 1933a; van Bemmelen 1949). The areal dimension of the original foundation limits horizontal expansion while subsidence may limit the vertical growth of the reef complex. The foundation of oceanic reefs is mainly volcanic in origin, thus their location is a product of volcanism and tectonic activity. For example, the Togian Ridge in the Gulf of Tomini is a product of volcanic activity, which is clearly demonstrated by the presence of exposed Neogene and young volcanic rocks on Batudaka and Waleabahi Islands.

O C E A N I C REEFS

595

Figure 12.9. Simplified schematic presentation of major oceanic reef types and their development on the Pacific Plate. After Scott and Rotondo

1983.

The development of barrier reefs and atolls in the Togian Islands occurred in response to tectonic activity, which caused subsidence as well as periodic uplifts of the volcanic foundation. Raised reef terraces on Batudaka Island, and elsewhere, are evidence of the complex tectonic setting of the region. Many oceanic reef systems in the archipelago have, however, developed on non-volcanic foundations. The atolls of the Tukang Besi group, off the south coast of southeast Sulawesi, apparently developed on top of a shallow undulating plateau, which during the Neogene was a part of a subsiding basin. During the end of the Tertiary, the Tukang Besi part of the basin was uplifted to a depth where coral reef development became possible. The subsequent subsidence during the Quaternary (van Bemmelen 1949) resulted in the formation of atolls. Oceanic volcanic islands are usually referred to as high islands. It is important to point out that, while many coral reefs associated with large high islands in the Pacific (e.g., Tahiti) are influenced by terrestrial processes, the effects are relatively minor when compared to continental landmasses, or large islands such as Java, Sumatra, Sulawesi or Borneo. Terrestrial processes on large islands (e.g., Sulawesi, etc.), while not necessarily detrimental to coral reef communities, may, however, have a significant effect on offshore reef development (fig. 12.10). Oceanic high islands in Indonesia vary from very large to very small, and may have a variety of reef types associated with them. For example, the active 1827-mhigh Siau in the Sangihe Archipelago is about 150 km 2 in area, while the neighboring 140-m-high Mahengetang has an area of only 0.25 km2 . Both volcanic islands are surrounded by extensive fringing reef systems, however, their coastlines are generally unexplored. Similar examples can be found in the southeast Banda Sea. Nila is a 781-m-high andesitic (pyroxene andesites) active volcano with an area of about 25 km , that last erupted in 1932. The volcano is surrounded by a series of

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Figure 12.10. Terrestrial processes on large islands, such as Sulawesi, may have a significant influence on the development of offshore "oceanic" reefs. A) The delta of Balinggara River, east Central Sulawesi, discharges large quantities of silt-laden runoff. The Togian Islands with their high diversity of reef types, ranging from fringing, barrier and atoll reefs, are located to the north. B) Terrestrial processes, such as erosion, have recently been greatly magnified by large-scale logging in the upper watershed. Photos by Tomas and Anmarie

Tomascik.

well-developed fringing reefs along both western and eastern slopes. However, barrier reefs have formed on the north and south sides of the volcano. The lagoon of the northern barrier reef is only about 2 km wide and 58 m deep. The eastern section of the barrier reef consists of a 1-km-wide reef that originates for about 2.5 km in a northerly direction, directly from the coastline (fig. 12.11). Coral communities on the west side of the barrier reef are very diverse with coral cover between 70%-90%. The shallow-water subtidal reef flat is dominated by massive Pontes spp. as well as Coscinaria exesa, which in particular forms large (up to 50 m ) monospecific stands. The seaward margin of the reef flat is a mixture of scleractinians and octocorals. The two most conspicuous alcyonarian genera are Sinularia and Sarcophyton. Another interesting barrier reef development occurred on the south coast of Nila, where a small barrier reef extends across the mouth of Solat Bay, thus enclosing a 32-m-deep lagoon. The coastline of the enclosed bay is lined with flourishing fringing reefs that extend up to 300 m offshore. The barrier reef across the bay is only 1 km long and about 250-m-wide. It seems that the small southern barrier reef may have developed either as an offshoot of a fringing reef, progressively expanding westwards until reaching the opposite shore. A more likely scenario is that the barrier grew on top of a submerged volcanic bank that initially blocked the mouth of the bay. Just seaward of the barrier reef is a large (500 m by 200 m) butterfly-shaped 40-m-deep blue hole, whose origin is not known.

Figure 12.11. Simplified morphological map of two small barrier reefs that developed on the north (A) and south (B) coasts of Nila Island. Nila is an active volcanic island in the south Banda Sea. Stippled areas indicate intertidal reef flats; dotted contours indicate boundaries of submerged reefs.

OCEANIC REEFS 597

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SEA-LEVEL FLUCTUATIONS Throughout geologic history, the growth (accretion), and therefore the size, shape and genesis, of coral reefs has been greatly influenced by the relative fluctuations of sea levels (Fagerstrom 1987). Indeed, it is now recognized that many present-day reefs are a product of reef growth cycles controlled by relative sea-level fluctuations throughout the Quaternary. Far from being globally uniform in their expression, however, sea-level fluctuations are a regional phenomenon that differs widely not only in the rate of rise and fall of the mean sea levels, but also in their general behaviour (Fairbridge 1961; Davies and Montaggioni 1985; Davies and Macintyre 1985). Global sea levels during the Holocene have experienced rapid changes with significant regional differences. These regional differences may be assessed, since both reef growth and the production of biolithofacies are sensitive to sea-level fluctuations (Davies and Montaggioni 1985). Major differences in sea level histories exist not only between the Caribbean and the southern Pacific and Australian regions, but also within the regions. Thus, while the wider Caribbean has experienced a continuous Holocene transgression, the behaviour of the sea-level curves between the northwest and southern Caribbean seem to differ. The northwest Caribbean sea-level curve is characterized by a rapid rise (5-6 mm.yr") prior to about 5000 yrs B.P., with a subsequent slow-down until the modern sea levels were reached some 2000 yrs B.P. (Neumann 1971; Lighty et al. 1982). In contrast, the sealevel curve for the southern Caribbean is characterized by a rapid rise (c. 7 mm.yr") to about 7000 yrs B.P.; a slower rise of c. 4 mm.yr" to within 1.6 m of sea level by 5200 yrs B.P., and a slow continued rise to modern levels (Davies and Montaggioni 1985). The sea-level curve for the Pacific is not only significantly different from the two Caribbean curves, but exhibits major regional differences as well. The South Pacific sea-level curve is characterized by a rise of about 6 mm.yr" until 6000 yrs B.P. when the present position was first reached, with a subsequent sea-level drop of about 0.9 m during the past 5000 years (Davies and Montaggioni 1985). The reconstruction of the Holocene sea-level curve for the Great Barrier Reef, based on information from 70 boreholes on 25 different reefs, revealed it to be a variant of the South Pacific sea-level curve (Davies et al. 1985). The sea-level curve for the Great Barrier Reef is characterized by a rapid rise (10 mm.yr"1) until the present sea level was reached between 5000-2000 yrs B.P. It is generally accepted that sea levels along the east coast of Australia have been relatively stable (i.e., stillstand) since about 6500 yrs B.P. Transgressions (i.e., rise in sea level) are caused either by classical Darwinian subsidence of the reefs foundation (i.e., volcanoes), which is tectonic in origin, or by an absolute sea-level rise (i.e., eustasy), or both. Considering the significant role of the antecedent topography on reef development during the Holocene transgression, the relative importance of tectonic and eustasy changes in sea level, on a geological time scale, needs to be placed in perspective. Fagerstrom (1987) argues that in the case of oceanic reefs, their prolonged development during the Cenozoic was influenced mainly by tectonism, with eustasy playing a subordinate role. However, the development and geomorphology of younger Neogene and Quaternary reefs, such as those on the tectonically stable and shallow Sunda Shelf, has been influenced mainly by eustasy. Local tectonic uplift rates, and thus regression, in the eastern regions of the archipelago are significant. Earlier studies of raised coral reef

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599

Figure 12.12. Quaternary sea-level history during the past one million years in the Indonesian Archipelago, deduced from study of uplifted coral reefs in Savu Sea (Hantoro 1992). From Hantoro

1995.

terraces on Atauro Island (Chappel and Veeh 1978), as well as more recent studies on Sumba (Pirazzoli et al. 1991) and Alor (Hantoro et al. 1994), offer a unique picture of Quaternary eustatic sea-level fluctuations preserved by the continual and rapid tectonic uplift during the past 50,000 years. Radiometric dating of uplifted coral reef terraces on Sumba Island indicates an average tectonic uplift rate of about 0.49 mm.yr" during the past one million years (Pirazzoli et al. 1991, 1993; Hantoro 1992). On Alor Island, radiometric dating of six major coral reef terraces, reaching an altitude of 700 m, suggests an average uplift rate of 1.0-1.2 mm.yr" since about 50,000 years (Hantoro et al. 1994). Based on the wealth of new information from the chrpnostratigraphic and neotectonic data from these significant studies, and in combination with the oxygen isotopic record, Hantoro (1995) was able to make a reconstruction of archipelagic sea-level fluctuations during the last one million years (fig. 12.12). Sea-level rise due to the subsidence of volcanoes, as envisaged by Darwin, is considered a local tectonic phenomenon. However, in a tectonically active region such as Indonesia, crustal warping, faulting, folding, subduction along active plate margins, and various other crustal movements may cause relative sea-level fluctuations, as well as lateral displacements on a much broader scale. It is, therefore, not surprising that the majority of coral reefs in the Indonesian Archipelago have been affected by both tectonic and eustatic changes. The raised coral reef terraces of Luwuk, Alor, Sumba, and Timor are excellent examples of tectonic forces at play. Based on their study of deep sediment cores, Netherwood and Wight (1993) have suggested that an interplay of tectonic and eustatic changes during the early Pliocene, in northeast Kalimantan, resulted in a sequence of four distinct reefgrowth cycles in an unusual delta-front setting. According to their interpretations

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of the logs, a relative sea-level fall, caused by a generation of a rollover anticline basinward of a large growth-fault system, resulted in sufficient shoaling to initiate reef development. Subsequent reef growth occurred under falling sea level conditions, as tectonic uplift was slightly outpacing early Pliocene eustatic rise (Netherwood and Wight 1993). Reef growth was terminated with rapid relative sea-level rise; in effect, the reefs were drowned. Whether the fast relative sea-level rise that terminated each growth cycle was caused primarily by tectonic (i.e., subsidence) or eustatic changes is not known. However, reef growth recommenced during subsequent tectonic pulses (Netherwood and Wight 1993). The cores clearly demonstrate that tectonism and eustasy have a considerable influence on the origin, location and subsequent development and growth of coral reefs. During the maximum sea-level low-stand, which occurred during the last glacial (Wisconsinian) (c. 20,000 yrs B.P.), the archipelagic sea levels were between 125 and 150 m below present levels (Hantoro 1995). Hantoro and Handayani (1993) estimated that approximately 2.9 x 106 km2 of the Sunda Shelf and 9.3 x 105 km2 of the Sahul Shelf were exposed during the last glacial maximum. The subaerially exposed Sunda Shelf was covered by a dense lowland tropical vegetation (Hantoro 1995), while the Sahul Shelf was covered by extensive eucalypt woodland open forest, grasslands-shrublands, with mangrove forests along the seaward margin of the shelf (van der Kaars 1991). Subsequent to the last glacial maximum, global sea levels continued to rise during the Holocene transgression (fig. 12.13). The present Java Sea is a product of the Holocene transgression, which was initiated some 11,000 yrs B.P. when the east Java Sea and the Sunda Strait were flooded (Park et al. 1992). Figure 12.14 provides a possible scenario of the Holocene transgression over the Sunda Shelf. The initial flooding of the Java Sea is expected to have occurred through the Molengraaff River system with the formation of large estuaries that were most likely diluted with fresh water. According to Brown (1991), the circulation pattern in the early Java Sea was characterized by a reversing-gyre pattern. Between 12,000 and 10,000 yrs B.P., the early Java Sea with its limited flushing and high influx of fresh water from Kalimantan and Java was nothing more than a large low-salinity western embayment of the Flores Sea. Most of the fresh water runoff from Sumatra entered the South China Sea, which at the time was separated from the proto Java Sea by topographic highs of Gelasa and Karimata Straits. It is likely that the earliest reef development in the present Java Sea may have occurred along the shores of Bawean Island, followed by reef development in Karimunjawa Islands. A drilling program in these islands would provide valuable information on coral reef development in the Java Sea during the initial stages of the Holocene transgression. At about 9500 yrs B.P., the westward-expanding Flores Sea connected with the Indian Ocean which penetrated eastwards through the Sunda Strait, thus significantly altering the oceanographic conditions of the Java Sea (Park et al. 1992). Subsequently, increased flushing through the Sunda Strait may have substantially increased salinities, which would have accelerated reef development at the western and eastern margins of the shelf. Brown (1991) suggested that coral reef development on Tunda Island (c. 10 km north of Tanjung Pontang), as well as Tijung and Pajung Islands (c. 10 km southwest of Thousand Islands), was initiated soon after the connection with the Indian Ocean was established some 9000 yrs B.P. Recent radiocarbon dating of a 7900-year-old coral fragment recovered from a depth of

SEA-LEVEL FLUCTUATIONS

601

Figure 12.13. Sea-level history during the past 300,000 years. I: last interglacial. G: last glacial. Modified from Chappel 1985 in Ongksoongo

1989.

19 m at Kepulauan Seribu (Pulau Putri Barat) adds some weight to this hypothesis (Park etal. 1992). The establishment of present-day oceanographic conditions in the Java Sea may date to about 8000 yrs B.P., when a new connection was forged with the South China Sea through the Gelasa and Karimata Straits between Sumatra and Borneo (East Kalimantan) (fig. 12.14). According to Brown (1991) and Park etal. (1992), initiation of carbonate buildup on the Kepulauan Seribu Platform may date to this event, when, more or less, the present-day monsoonal circulation was established thus giving the Java Sea its fully marine character. This is supported by data obtained from a core taken on Pulau Putri Barat, where at a depth of 28 m (representing early Holocene flooding) the sediment composition consisted of terrigenous clay and carbonate fine sediments characteristic of shallow-water protected coastal environments similar to those found today along the north coast of Java (Parketal. 1992). Whether the rates of sea-level rise during the Holocene transgression were globally uniform is not known, mainly because of local and regional differences in tectonic settings. Nonetheless, there is an agreement that during the Holocene transgression, between 18,000-6000 yrs B.P., relative sea levels rose at rates between 5 to 10 mm.yr" (Park et al. 1992). As the sea levels rose, many deep-water reefs, especially those in marginal environments with shallow euphotic zones (i.e., turbid shelf waters), drowned. Reefs drowned, because they could not keep up with rapidly rising Holocene sea levels that displaced the euphotic zone upwards. The relative downward displacement of coral communities into the aphotic zone effectively terminated subsequent reef growth, since coral growth rates (thus reef growth) are dependent on zooxanthellae photosynthesis. However, most shallowwater reefs are presumed to have kept pace with the sea level as it continued to rise throughout the Holocene. The youngest reefs on the Sunda Shelf were estab-

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Figure 12.14. Holocene transgression over the Sunda Shelf from 12,000 to 6000 yrs

B.P.

Modified from Brown 1991.

SEA-LEVEL F L U C T U A T I O N S

603

Figure 12.14. (Continued.)

lished during the Holocene transgression on suitable new foundations such as the seaward margin of the shelf (e.g., Great Sunda Barrier Reef), along the coastlines of its numerous continental islands (e.g., Belitung, Karimunjawa), and on topographic highs of the shelf (e.g., Kepulauan Seribu). The basement of the Kepulauan Seribu Platform consists of stiff marine mudstone of apparent Pleistocene age (Park et al. 1992). Based on their interpretation of core data, Park et al. (1992) speculate that the 25-30 m reefal build-up on the Seribu Platform apparendy occurred on minor topographic features, which were either inherited from the Pleistocene or were abrasional products of early Holocene storm and tidal currents. Sea levels in Southeast Asia reached their present levels about 6500 years ago and continued to rise until 4500 years ago when the Holocene sea-level high-stand was reached, somewhere between 1-4 m above the present-day mean sea level (figs. 12.15 and 12.16). The Holocene maximum corresponds to the period when most of Arctic and Antarctic ice-sheet melt-waters were added to the oceans (Lambeck and Nakada 1985). Since the Holocene high-stand (i.e., the past 4500 years), coral reefs in Southeast Asia have been experiencing the effects of gradual sea-level fall

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Figure 12.15. Sea-level history of Southeast Asia during the past 8000 years. Modified from Brown 1991.

Figure 12.16. Sea-level history of Malacca Strait during the past 8000 years. Source: Geyeth et al. 1979.

REEF GROWTH AND DEVELOPMENT

605

Figure 12.17. Sea-level history for west and central Indonesia during the past 8000 years. From Hantoro

1995.

(i.e., regression). Holocene sea-level reconstructions from west and central Indonesia support this generalization (Hantoro 1995) (fig. 12.17). Additional support for this hypothesis comes from recent radiocarbon dating of coral samples from Pulau Putri Barat (Kepulauan Seribu) recovered from about 1.0 m above the present-day mean sea level and dated to 4400 yrs B.P. (Park et al. 1992). Thus, the archipelagic sea-level curve differs significantly from both the Caribbean and Pacific reconstructions. It is also clear that tectonism in many regions of the archipelago has modified the eustatic effects. The eastern parts of the archipelago, in particular, have experienced rapid and relatively frequent tectonic uplifts.

REEF G R O W T H A N D DEVELOPMENT According to the classical Darwinian subsidence model, coral reefs grow mainly upwards as they track the relative sea-level rise caused by the continuous subsidence of their foundation. However, it should be noted that reef growth, or more properly reef accretion, does not proceed in an upward direction only, since once stable sea levels are reached coral reefs will expand in a lateral direction as well. Reef growth is a complex phenomenon resulting from a continual interplay between constructive and destructive processes, both biotic and abiotic. As Fagerstrom (1987) points out: "For reefs to achieve positive topographic relief, their long-term upward accretion ("growth") rate must exceed the accumulation rate of sediments on the adjacent level-bottom". Reef growth or accretion can be measured either as a linear extension in vertical or lateral directions, or by the net production (deposition minus removal) of

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CaC0 3 per unit area. Within this framework, the constructional role of reefbuilding corals and coralline algae, as well as a host of other reef-building organisms (e.g., foraminiferans, some sponges, bryozoans, molluscs, echinoids), is reinforced by sedimentary depositional processes operating continually in a variety of reefal environments. Reef construction and growth (i.e., positive processes) includes a number of interrelated processes that include: 1) new reef build-up through the growth of numerous reef-building organisms upon the original framework; 2) binding and cementation; 3) baffling; 4) sediment deposition and accumulation; and 5) encrustation (i.e., protective coating). Note that reef processes that maintain and consolidate the reef structure are lumped with the constructional processes (apologies to geologists). Reef Accretion Rates Growth of coral reefs varies along both spatial and temporal scales. Upward growth rates of most reefs are in the range of 1-15 mm.yf , while lateral extension rates are seldom above 2.5 mm.yr"1 (Fagerstrom 1987). During the early Holocene, the forereef facies tracked the sea-level rise (in excess of 4 mm.yr") with reef accumulation rates of about 5-18 mm.yr" (Kayanne et al. 1993). Growth rates of reefs in the archipelago follow similar patterns. For example, reefs in the Kepulauan Seribu Complex grew at a rate of 5-10 mm.yr" from 8000 yrs B.P. to 4500 yrs B.P. Since 4500 yrs B.P., sea levels in this region have been relatively stable and reef growth has been largely outwards, leading to the coalescence and enlargement of the build-ups (Park et al. 1993). It should be kept in mind, however, that not all parts of a reef complex grow at the same rate. The three reef zones where CaCO s production usually exceeds the rate of destruction are the algal ridge, reef crest and the upper forereef slope. The algal ridge zone is made of coralline red algae and forms at the seaward margin of the reef flat, within the intertidal to a depth of about 2 m or less. The algal ridge is the highest energy zone of a reef complex. There are 10 genera of coralline algae (with 5-15 species) that flourish in this environment (Davies and Montaggioni 1985). The dominant genera of coralline algae in the Indo-Pacific are Neogoniolithon, Porolithon, and Lithoporella. In Indonesia, coral reefs with well-developed algal ridge zones are restricted mainly to the southern coastline of the Sunda Arc, and to the high-energy regions along the north margin of the archipelago. Kuenen (1933a) commented on the conspicuous absence of the so-called Lithothamnium ridge from the reefs in the archipelago. Note that the term Lithothamnium ridge has been replaced by algal ridge. Lithothamnium ridge was a misleading term since the dominant group of coralline algae responsible for the actual carbonate deposition in this high-energy environment are species of the genus Porolithon. The itinerary of the Snellius Expedition did not include stops along the southern coastlines of Java and Sumatra, thus Kuenen (1933a) was not able to observe the algal ridges present on most of the fringing reefs along the Indian Ocean coastline. Kadi (1988) conducted a survey of the coralline algae along the south coast of Lombok and found 11 species on the seaward margin of the reef flats. The dominant corallines on the outer reef flat were Porolithon spp., Lithothamnium calcareum, L. asperulum, L. mesomerphum and Goniolithon strictum. Studies conducted in different regions of the Pacific show that while accretion of algal ridges in the Mariana Islands proceeds

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at a rate of 0.5-1.2 mm.yr" (Kayanne et al. 1993), it is only 0.3 mm.yr"1 in Hanauma Bay, Hawaii (Easton and Olson 1976), but up to 2 mm.yr"1 on the Great Barrier Reef (Davies and Hopley 1983). Located just seaward of the algal ridge is the reef crest. The reef crest is a high-energy zone which forms a seaward rise at depths above -5 m and below lowtide level. It is an actively growing zone dominated by corals adapted to highenergy environments such as Acropora humilis, A. palifera and A. gemmifera. Accretion rates on the reef crest are 1.0-3.2 mm.yr"1 for the Mariana Islands (Kayanne et al. 1993), 2.9 mm.yr"1 for Hawaii (Hanauma Bay) (Easton and Olson 1976), and 1.3 to 4.2 mm.yr" in the eastern Pacific (Glynn and Macintyre 1977). The fore-reef slope is the most vigorous reef zone, consisting of high-diversity coral assemblages and unconsolidated bioclastics that form a thick reefal sedimentary unit. The fore-reef slope is a moderate- to low-energy environment usually below 3 m depth. It is characterized by a catch-up style of reef development with measured accretion rates of 5-16 mm.yr"1 in the Mariana Islands (Kayanne et al. 1993), 5 mm.yr" in the Palau Islands (Easton and Ku 1980) and 5-18 mm.yr" on the Great Barrier Reef (Davies and Hopley 1983). The range of coral reef accretion rates are summarized in table 12.1. Comparable values for Indonesian reefs are not available. However, it is expected that vertical growth rates of submerged oceanic platform reefs in the upwelling regions of the Banda Sea may be relatively high, based on our observation of the vigorous Aopjfrora-dominated (e.g., Acropora palifera, A. brueggemanni, A. monticulosa) reef flats, reef crests and upper reef slopes. Reef Development Environmental conditions in tropical coral seas do not always favour initiation of reef development, or continual accretion of existing reefs. As discussed earlier, relative sea-level fluctuations, due either to eustasy or diastrophism or both, have considerable influence on the ultimate fate of many coral reefs. Understanding how coral reefs respond to sea-level fluctuations is especially important with regards to the Holocene marine transgression, which to a large extent shaped the geomorphology of many present-day reefs. Reef development on the Sunda and Sahul Shelves is a direct result of the Holocene transgression, which in effect opened up vast new areas for reef colonization. It is most likely that Holocene reef development on the Sunda Shelf occurred against or over topographic highs that were

Table 12.1. Examples of Holocene accretion rate of major coral reef zones. Reef zone

Biotic components

Algal ridge Reef crest

Calcareous algae Porolithon spp. Scleractinian Corals Acropora spp. Mixed community; massive dominate

Reef slope After Kayanne 1993.

Depth (m) 0-2 LLW - 5