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The evolution an d biogeography of Neogene M icronesian O stracodes: The role of sea level, geography, and dispersal R oss, R o b e rt M errill, P h .D . Harvard University, 1990

Copyright © 1990 by Ross, Robert Merrill. All rights reserved.

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H A R V A R D U N IV E R S IT Y THE GRADUATE SCHOOL OF ARTS AND SCIENCES

TH ESIS A C CEPTA N CE C E R TIFIC A T E (T o be placed in Original Copy) T h e undersigned, appointed by the Division D epartm ent o f Earth and P l a n e t a r y S c i e n c e s Com m ittee

have examined a thesis entitled The E v o l u t i o n and Biogeograpl o f M i c r o n e s i a n O st r a c o d e s : I m p l i c a t i o n s f o r t h e Role o f Geography and Cl i mat e Change

presented by

Robert M. Ross

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Thomas Trqni.n

Date

May 15, 1990

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The Evolution and Biogeography of Neogene Micronesian Ostracodes: The Role of Sea Level, Geography, and Dispersal

A thesis presented by Robert Merrill Ross to The Department of Earth and P’^netary Sciences in partial fulfillment of the requirements for the degree of Doctor of Philosophy in the subject of Geology Harvard University Cambridge, Massachusetts June, 1990

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© 1990 by Robert Merrill Ross All rights reserved.

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Abstract

Ostracodes from the Neogene of Enewetak, m odem Micronesia, and other Pacific localities were studied to understand better the evolution of marine organisms on tropical oceanic islands. Temporo-spatial patterns in ostracode species composition and species richness, and in the morphology of species of the genus Loxoconcha, indicate that the evolution and biogeography of Quaternary Micronesian benthic ostracodes has been controlled by local extinction due to high amplitude glacio-eustatic sea level fluctuations and distance among atoll lagoons. Hierarchical clustering in species composition suggests that colonization takes place primarily from neighboring lagoons. Change from Late Pliocene to Holocene assemblages at Enewetak is concentrated at disconformities rather than within zones between them, suggesting that survival of sea level drops or ability to recolonize determine the character of species assemblages. Assemblage turnover from the Late Miocene to Early Pliocene, during relatively stable high sea level, is more gradual in nature, and the turnover trend is nearly independent of the Late Pliocene to Quaternary trend toward m odem samples. Among modem Micronesian lagoons, the highest species richness occurs on high islands to the west; similar diversity is associated with high island conditions in the Miocene at Enewetak. Species richness drops to levels typical of m odem lagoons over a disconformity near the Plio-Pleistocene boundary. Morphological variation within loxoconchid species is lower among modern lagoons across Micronesia than through geologic time at Enewetak. Possibly variants first colonize one Micronesian lagoon from outside the region, and then spread throughout Micronesia. Most temporal morphologic variation is among temporally stable variants that go locally extinct primarily at disconformity boundaries. In most species there is little net trend in morphology since the Miocene. Morphologic and assemblage data provide complementary information on evolution and biogeography in the region. Both suggest that populations at any one point in time, even without gene flow, are more related to geographically surrounding

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populations than to populations that have preceded them in geologic time at the same lagoon. Extinction and rapid recolonization have therefore acted to create temporal heterogeneity and spatial homogeneity within linages and perhaps within assemblages in Micronesia.

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Table of Contents

Abstract..................................................................................................................................iii Acknowledgments.................................................................................................................vi Chapter 1: Introduction......................................................................................................... 1 Historical Context: Geology and Paleoceanography of the W estern Equatorial Pacific............................................................................................. 6 Ecologic Context: Ecology, physiography, and oceanography of Micronesian lagoons................................................................................................. 20 Summary................................................................................................................................ 30 Chapter 2: The role of sea level and geography in the species Composition of Micronesian ostracode assemblages...................................................38 Methods............................................................................................................................... 43 Patterns................................................................................................................................ 46 Discussion........................................................................................................................... 59 Conclusions......................................................................................................................... 69 Chapter 3: Taxonomy of Micronesian loxoconchids.....................................................103 Summary comments........................................................................................................ 146 Chapter 4: The role of frequent local extinction and recolonization in the evolution of Micronesian ostracodes..................................... 257 Material............................................................................................................................. 261 Measurements...................................................................................................................263 Results............................................................................................................................... 270 Discussion......................................................................................................................... 285 Conclusions...................................................................................................................... 294 Chapter 5: A final overview............................................................................................ 365 Bibliography........................................................................................................................ 372 Appendix 1: List of modern samples.............................................................................393 Appendix 2: Species presence/absence and abundance data..................................... 405 Appendix 3: Loxoconchid morphologic measurement data........................................ 507

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Acknowledgments

Looking back, it is amazing how many people contributed in some way to this project. I cannot really ever hope to repay Thomas Cronin, advisor, boss, friend, and running partner; Tom first suggested this study, did much of the much of the pioneering work that m ade it possible, and freely donated data, specimens, ideas, equipment, support, advice, beer, and just about anything else you could think of. I also want to add a special thank you to Tom's other (better?) half, Margarita, who not only helped supply additional support, friendship, and beer, but also m any excellent meals to my otherwise subsistence-level diet. Nor can I hope to repay Stephen Gould, who has shaped my thinking more than I care to admit, and who not only has supported me in all ways possible, but who also wrote the letter to Tom Cronin that got this whole thing started. I never got to the forams, Steve, but they're picked and on my list of things to do. I also want to thank Andy Knoll, who endured two frantically written term-papers nearly as long as this thesis, and who never laughed outright at my periodic early faux pas. Thanks for putting me up in Canberra too. And also a special thanks to Peter Williamson, who took my thesis proposal home and rewrote it, and who provided unabashed enthusiasm and support. Peter McCall told me to go to the graduate school with the best students, because they would be the ones from whom I would learn the most. Great advice, but I'm sure I never learned as much from any student as I learned from Peter himself. I met Peter while visiting the Case Western campus in high school, but little did I know that he would guide me through the equivalent of 7 paleo- courses, a summer of research, and several evenings at the Euclid tavern. So I give an unconventional thank you to a pre-thesis advisor. But about those students.

Warren Allmon guided my survival during my years

in Cambridge, and I doubtlessly prolonged his own stay in Cambridge through my endless distracting trips up and down the first floor MCZ corridor. Sorry Warren.

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Emily CoBabe provided meals and helped me retain my sanity in late stages of this project. Greg Mayer patiently attempted to rectify my naive biogeographical and reviewed the first two chapters of this thesis. I would like to thank in particular my field assistants, who braved denizens of the deep, unknown disease, political coups, and other hardships to collect animals they could not see. Thanks to Consuelo Marino, who was the ultimate field trooper. Thanks to Tula Economou, who helped me get through third world sites without so much as an upset stomach. And thanks to Mandana Azar, whose traveling aplomb made third world travel easy. Lest these acknowledgements become another chapter of this thesis, I cannot unfortunately devote a paragraph to every thank you. All of the following have contributed to this thesis through help, discussion, encouragement, and friendship: Robin Aiello, Felice Apter, David Backus, Matthew Cotton, Holly Dale, Harry Dowsett, Jon Ellman, Ronald Eng, Felicita D'Escrivan (I can't find my keys), Michael Foote, Thomas Gibson, Stephen Grant, James Hamilton, Timothy Heaton, David Kendrick, Craig Mello, Paul Morris, Richard Norris, Gustav Paulay, Agnes Pilot, Jane Rose, Cindy Schneider, and Mark Schramm (on whose computer this is written). I give a particular thank you to Robert Blodgett, who gave m e a crash course in darkroom skills and personally saw to it that my plates got made. Though I cannot hope to thank them all here, I would also like to thank the many people who loaned samples, helped me collect samples, or provided field support for this study. Financial support for this study was provided by grants from the American Association of Petroleum Geologists, American Museum, Geological Society of America, Harvard University Department of Earth and Planetary Sciences, National Science Foundation, Paleontological Society, Sigma Xi, and Stephen Gould. The U.S. Geological Survey, Branch of Paleontology and Stratigraphy, Reston, Va provided facilities during my employment there from the fall of 1988 to the spring of 1990. I would like to provide a very special thank you to my family, who have been so supportive over such a long period without a real job.

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Introduction

In 1881 Charles Darwin wrote to Alexander Agassiz: "I wish that some doubly rich millionaire would take it into his head to have borings m ade in some of the Pacific and Indian atolls..." (Darwin 1881). Darwin hoped such cores would contain a thick shallow marine limestone cap over a volcanic edifice, which would confirm his theory that atolls are formed by upw ard reef growth over a volcano on a subsiding sea floor. Seventy-three years later, the U.S. Geological Survey became the first to core the basalt basement of an atoll, a t Enewetak in the Marshall Islands (Emery et al 1954, Rosen 1982), proving Darwin right over 100 years after he formulated his idea. Many cores have been taken at Enewetak since, the most recent set in 1985 (Henry et al 1986), forming one of the most thoroughly studied long-term records of shallow marine sedimentation in existence. The cores contain a well-preserved fossil record. Darwin would have been pleased to know that the very sediments confirming his first and most encompassing geological theory also provided excellent material for studies of the other realm for which he is more generally known — evolution. In this study I establish temporo-spatial patterns of morphologic and assemblage evolution in Micronesian island ostracodes, using fossil material from Enewetak Atoll and m odem material from other islands across Micronesia. My primary goal is to provide an understanding of the effect of changes in Neogene sea level and oceanic island settings on lagoonal marine biotas and on the nature of morphologic evolutionary patterns in general. "Oceanic" islands are islands fcrmc d de novo from the oceanic floor by volcanic activity and are separated from continental land by deep water. Oceanic island biotas have been of considerable interest since Darwin recognized them as "laboratories" for the study of ecology and evolution (Carlquist 1974, Brown and Gibson 1983) -

islands provide numerous replicated natural experiments on

landmasses small enough to be characterized ecologically and taxonomically. The best known aspect of island biotas is adaptive radiation among terrestrial animals, such as Darwin's finches (Geospiza) in the Galapagos islands, Partula in Polynesia, Drosophila and

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honeycreepers in the Hawaiian islands, and Cerion and Anolis in the West Indies. The enormous am ount of literature generated about these biotas has significantly influenced our perception of evolutionary processes. More recently islands have been recognized as areas in which to test ecological theory and to integrate ecology and biogeography. Most of this research has revolved around testing theories of island biogeography (Case and Cody 1987), particularly the equilibrium theory of MacArthur and Wilson (1963, 1967); these studies have also concentrated upon terrestrial organisms. Though always controversial, island biogeographical theory in general, and MacArthur and Wilson's theory in particular, have resulted in two decades of model testing, and have been applied at every scale from parasitic colonization to Phanerozoic global marine diversity. In contrast, the study of marine organisms has until recently played a relatively minor role in the study of island evolution or ecological biogeographic theory. While the ocean does not occur in discrete island patches, small shallow water areas (such as oceanic island lagoons) separated by wide expanses of deep water form discrete islands of habitation for organisms adapted to shallow environments. The lack of island biogeographic study of tropical marine island biotas at first seems surprising in light of the fact that these biotas contain the world's most spectacular diversity and disparity1 of living marine organisms. However, individual species tend to be widely distributed and island biotas (oceanic islands in particular) are rarely home to clusters of endemic species (but see Pokomy 1969). While dispersability of marine organisms makes them potential subjects for the subject of equilibrium island biogeographic models, the ecological controls are much less well understood for marine environments. Marine zoogeographers have primarily concentrated upon issues at larger scales, such as the place of origin and maintenance of such diversity gradients and the role oceanic islands have played (e.g. Ladd 1960, Briggs 1974, 1981, Kay 1985, Kay and Palumbi 1987, Stehli and Wells 1971, Rosen 1984, Hallock 1987). Recently, there has been substantial interest

'Disparity here refers to diversity at high taxonomic levels.

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in the documentation of temporal changes in vicariant boundaries (McCoy and Heack 1976, 1983, Heck and McCoy 1978, 1979, Rosen 1984, Vermeij 1989) and the use of areas of endemism (Springer 1982, Rosen 1984) to understand large-scale biogeographic relationships. Some authors have further suggested that macroevolutionary trends may have been controlled in part by the distribution of tropical oceanic islands or by largescale changes in tropical reef systems in general (Fagerstrom 1983, 1987, Rosen 1984, Jablonski 1985, 1986, Jablonski and Flessa 1986, Hallock 1987, Hallock and Schlanger 1986, Vermeij 1989). The resolution of these debates will have an important influence on our view of the origin of marine biotic distribution and diversity patterns. However, very little detailed data has been available on what happens to organisms over geologic time at oceanic islands, or on their distribution and morphologic variation among island within an archipelago. It is these gaps in our knowledge that I will attem pt to begin to fill in the chapters that follow, using modem and fossil benthic ostracodes from Micronesia. This exploration is divided among three chapter (Chapters 2 to 4). Chapter 2 is devoted to temporo-spatial patterns of species composition and species-richness, studying especially the impact of Plio-Pleistocene sea level fluctuations upon the evolution of the biota. Chapters 3 and 4 are devoted to the evolution of loxoconchid ostracodes, chapter 3 on taxonomy and chapter 4 on quantification of temporo-spatial patterns of morphologic variation and evolution. The patterns, I hope, will be of interest not only within the context of island biogeography, but also for further documentation of long-term species-level evolution. In chapter 5, I summarize the results, and suggest the two sorts of data are concordant and together provide more insight than could be had using one alone. I will briefly speculate on the potential role of oceanic islands and tropical reefs in general for larger-scale evolutionary patterns.

The ostracode animal

Ostracodes are small bivalved crustaceans, generally .5-2 mm in length as adults. Plate 1 illustrates several ostracode carapaces. The calcareous valves are hinged along

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the dorsal margin (top of the specimens in Figures 1, 2, and 3). Ostracodes are favorable material for evolutionary and biogeographic studies, especially those that can take advantage of cored fossil sediments. Ostracodes' small but robust carapaces are often present in large numbers in m ost aqueous sediments. Their carapace mirrors aspects of the internal anatomy and epidermal surface of the ostracode body, and thus provides detailed for the study of taxonomy and evolutionary history. These aspects are described in more detail in later chapters. Ostracodes are believed to disperse between oceanic islands by attachment to floating algae (Teeter 1975). Ostracodes have no known planktonic larvae. Unlike lacustrine ostracodes, marine ostracodes are not known to possess a desiccation resistant egg that might enable above-water dispersal on birds feet or by wind. Other organisms of their size, such as benthic foraminifera and amphipods, are similarly widespread despite lack of special dispersal mechanisms. The remainder of this chapter provides the geologic, oceanographic, and ecologic context of this study. Each of the following chapters draw s from information presented in the introduction.

Geography

"[The Indo-Pacific is]...the greatest of all marine regions. I regard it as constituting but a single division, because all explorations go to show that throughout its area there is a basis, as it were, of identical species, giving a uniform character to its fauna. It is the realm of reef-building corals and of the wondrously beautiful assemblages of animals, Vertebrate and Invertebrate that live among them, or prey upon them.” Edward Forbes 1856

The equatorial Indo-Pacific comprises 30% of the Earth's surface and wraps around

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two-thirds of the earth's circumference. Within this vast region, just to the north and east of its center, lies Micronesia, the region whose biota I will concentrate upon in this study. Micronesia is approximately the size of the continental United States in over-all area, but only slightly larger than Rhode Island in combined land area. The islands number about 2400 (Bryan 1961). The biota and geology of Micronesia have much in common with other parts of the oceanic Pacific and the greatei Indo-Pacific region, and it is my hope that studies of Micronesia will have much to offer to the study of tropical islands in general. Thus I will frequently refer not only to Micronesia but to other areas within the expanse of the equatorial Indo-Pacific. Micronesia is variously defined, depending on the purposes and field of the author. Most commonly Micronesia is defined to include the Marianas, Caroline, Marshall, and Gilbert Island archipelagos, and I will follow this convention. The Marianas, Carolines, and Marshalls form a natural geographic unit of island chains north of the Equator; they are connected to the next closest set of oceanic islands through the Gilbert Islands, which run across the equator southeast toward Polynesia (Figure 1-1). Ethnological and biogeographic patterns tend to reflect this geography and thus ethnologists and biogeographers often define Micronesia in the same way. After centuries of colonization, the peoples of Micronesia recently became more politically independent. The people of the Gilbert Islands, formerly under British rule, joined with those in the Line Islands, Phoenix Islands, and Banaba (islands usually placed within Polynesia) and named their nation Kiribati.

The Marianas, Carolines, and Marshall

Islands, put under the trusteeship of the United States after the second World War, now comprise several nations. The Marianas are split into two politically distinct regions: the southernmost and largest island, Guam, which is a Territory of the United States, and the North Marianas, which are a United States commonwealth. The Caroline Islands have become the Republic of Palau (=Belau) and the Federated States of Micronesia. The "FSM" is divided into 5 states running west to east, each of which is named after a prominent high island that serves as the capital. The Marshall Islands are the Republic of the Marshall Islands, with the capitol at Majuro.

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HISTORICAL CONTEXT: GEOLOGY AND PALEOCEANOGRAPHY OF THE WESTERN EQUATORIAL PACIFIC Geology Indo-Pacific geology

The geology of the Indo-Pacific, particularly the central Indo-Malaysian region, is extremely diverse and complex. The geology of Southeast Asia has been reviewed by Hutchinson (1988). Much of the complexity surrounds the collision of three large plates, the Pacific, Australian, and Eurasian plates, and a 4th smaller Philippine plate. A triple junction (or possibly two, depending on the interpretation of the southwest portion of the Philippine plate boundary) of colliding plates occurs near the northwest tip of New Guinea. The ensuing earthquakes and active volcanism along the volcanic arcs makes the Indo-Malaysian region one of the most tectonically active areas in the world. Volcanism has contributed substantially to the land area, especially along the continental edges of the plates. The Pacific plate is the one of the largest plates on the Earth and the only major plate without a substantial continental land mass. Oceanic islands and submerged sea mounts cover the floor of the plate. Many volcanoes dot the southwest edge of the plate at subduction boundaries. The pattern of subduction along the Pacific plate margin, which reverses direction at several points and has not been detected at others, is not entirely understood. The saw-tooth configuration of island arc chains may be created by the retarded motion of plates at subduction boundaries where unusual volumes of lithosphere (such as migrating volcanic islands or fragments of continental crust) are being consumed (Scoffin and Dixon 1983). The majority of the volcanoes on the Pacific plate are formed on the inner part of the plate. Most of these are submerged below sea level. Some of the intraplate volcanoes are arranged in linear trends and are believed to have been formed by stationary "hot-spots" in the mantle, below the plate, that periodically erupted (typically) tholeiitic basalts (Wilson 1963, McDougall 1971, Molnar and Atwater 1983, Clague and

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Jarrard 1973, Jarrard and Q ague 1977, McDougal and Duncan 1980, Duncan and Clague 1985). Most of the Pacific islands "chains" however, have much scatter around their linear trend, and do not easily fit a single stationaiy hot-spot model (look at the scatter, for example, within the Carolines and Marshalls-Gilberts, Figure 1-1). In addition, the foundations of some chains, such as those of the Marshalls and Line-Tuamotus, seem to have been simultaneously active, rather than sequentially active (Q ague 1981, Schlanger 1981).

Micronesian geology

"All of the Marshalls and Gilberts are coral atolls or islands, while the Marianas are volcanic. Micronesia's only active volcanoes erupt in the Northern Marianas. The Caroline Islands include both volcanic and coral types. Yap and Belau are the sedimentary tips of an undersea ridge stretching between Japan and New Guinea, thus their structure is continental. This surprising variety of landforms makes Micronesia a geologist's paradise." David Stanley (1985)

Micronesia consists of both island arcs and intraplate volcanos (that may be partially or entirely hot spot islands) of varying age. The Marianas Islands are an island arc along the eastern margin of the Philippine plate where ’he Pacific plate under-rides the Philippine plate. The well-known Marianas trench, the deepest point in the world's oceans, occurs directly to the east. Guam belongs to the western half of the Western Marianas Ridge (Bloomer et al 1989). The geology of Guam was summarized by Tracey et al (1964). The island was apparently formed by the Eocene and experienced explosive volcanism again in the Miocene; Eocene reef facies are present immediately above the earliest volcanics. There was a period of emergence in the Miocene during which no carbonates were deposited,

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but carbonate sedimentation resumed in the Plio-Pleistocene. The island has apparently been uplifted several meters since the Pleistocene (Curray e t al 1970). Some of the Caroline Islands east of Yap and Palau are a relatively young chain resulting from hot spot activity that has ceased within the past 1 my (Keating et al 1984a). Keating et al (1984b) reported that Truk was initially formed about 10.9 Ma at 4.4 degrees north of the equator (it is now 7.3 degrees north), and that it received posterosional lavas at 4.8 Ma. They suggest that the main shield-building lavas of Pohnpei were erupted about 5.2 Ma, with deposition of post-erosional lavas within the past 1 my. Kusaie, the eastern-most Caroline Island, was formed only 1 Ma. Because Truk, Pohnpei, and Kusaie are progressively smaller with more alkaline lavas they argue that the Caroline hot spot is declining in activity. In contrast, the Marshall Islands are among the oldest known sea floor volcanic structures with emergent land. A voluminous literature dealing with the origin and geologic history (as well as many other aspects of natural history) of the Northern Marshalls has developed since the atomic bomb tests commenced in 1947 [Ladd (1973) gave a brief review of the literature up to that time; W ardlaw and H enry (1986) reviewed the history of nuclear testing and coring at Enewetak; Devaney et al (1987) edited two volumes summarizing the natural history, which include a paper on the geology by Risvet]. The region around the northern Marshall islands is believed to have experienced considerable volcanism by the Late Cretaceous (Schlanger et al 1987), as evidenced by Late Cretaceous ages of several submerged volcanic edifices (Hein et al, 1988 and numerous submarine flows and sills discovered at DSDP site 462 in the Nauru Basin (Schlanger and Moberly 1985). Kulp (1963) dated volcanics immediately below the oldest limestone cap at Enewetak at 50 and 59 million years using K /A r dating. A phosphatized Globogerina ooze from cracks in a tuff-breccia near the surface of Sylvania Guyot, adjacent to Bikini, was Early Eocene (Ladd 1973); Schlanger et al (1987) suggested on additional geophysical and geologic evidence that Sylvania Guyot and Harrie Guyot (south of Majuro) are drowned atolls with volcanic edifices that apparently formed coevally with those of Enewetak and Bikini. No radiometric age data for the Gilbert chain have been reported, but according to

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Morgan (1972) the azim uth of the Gilberts and Marshalls suggest it is contemporaneous with that of the Line Islands and Emperor Sch Mounts, which mugc frcia 93 to 44 and 64.7 to 42 mya respectively (Duncan and Clague 1985). However, present evidence does not suggest a linear arrangement of ages for these volcanoes. Some authors, as an alternative to hot spot formation, have suggested reviving M enard's postulated Darwin Rise (Schlanger 1981, Rosen 1982), a large volcanic edifice which Menard (1964) believed had gradually subsided.

Oceanic island reef formation

To understand the fashion in which tropical lagoon biotas have evolved, it is necessary to understand how these lagoons developed, how they m ight be affected by sea level fluctuations, and how different parts of Micronesia m ay have been differentially affected. The most general tenets of Darwin's (1842, 1874) atoll theory, later modified by m odem plate tectonic theory, explain a great deal of the variance in the broad-scale pattern of high-islands and coral atolls in the Indo-Pacific. Darwin proposed that reefs growing along the sides of a volcano ("fringing reefs") will grow upw ard to maintain an equilibrium distance from the sea surface, as the volcanic edifice subsides beneath the sea. As the part of the volcano still emergent becomes smaller, a lagoon opens between the reef ("barrier reef") and volcano. Once the volcanic landmass has disappeared completely beneath the sea, the structure is called an "atoll". In Darwin's time the primary objection to his theory was that no mechanism was known to create subsidence over the breadth of the Pacific sea floor. It has only been within the past two decades that sea floor subsidence has been attributed to heat anomalies in the aesthenosphere. It is well-known, however, that many exceptions to the "rules" exist and that, even for atolls that behave well under the general subsidence model, temporary historical deviations are commonplace. The history of theories of the formation of barrier reefs and atolls published after the Darwin (1842) are reviewed by Steers and Stoddart (1977), and the major developments in the understanding of reef structure and distribution over the past two decades are reviewed by Hopley (1982)

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and Scoffin and Dixon (1983). Daly (1910) and many followers provided the most serious challenge to subsidence theory in suggesting that the series of reef morphologies, from fringing reef to atoll, could be explained better by upgrow th of reefs along the rims of landmasses that had been erosionally planed during glacial sea level drops. Purdy (1974) reviewed favorable evidence for the influence of sea level, but suggested that in fact karstic topography rather than erosional planing played the major role in m odem reef development. It is now indisputable that both eustatic sea level fluctuation and long-term subsidence play important roles, but only recently have researchers begun to model reef formation using both factors (Rosen 1982); the influence of differing temporal scales and rates, and the fact that each results in superficially similar geomorphologies makes the influences difficult to decouple. It is unknown how atoll morphology developed over intervals of the Cenozoic prior to large-scale rapid large-scale sea level change, during which time reef-derived lagoonal sedimentation probably outstripped volcanic subsidence. Presumably no m odem analogues exist because no reef area would have been completely isolated from Quaternary climate change. Therefore interpretations of the nature of Late Miocene atolls m ust depend upon modeling and detailed geologic investigations of relatively well known areas such as Enewetak and the Great Barrier Reef. Models of reef morphology relying on eustatic sea level fluctuations and a gradually subsiding volcanic edifice fail to explain a number of aspects of mid-Pacific geology. Deviations from these models are manifested in uplifted islands within the central Pacific and in islands that appear to have subsided too rapidly to fit the model of sea floor subsidence suggested by Parson and Sclater (1977) (Detrick and Cough 1978, Schlanger et al 1987). Woodroffe (1988) reviewed mechanisms that may result in relative uplift of fossil reefs; a number of these refer especially to islands along plate boundaries, but others are more generally applicable. He divided these mechanisms into 3 categories, including local tectonic uplift, regional isostatic compensation or lithospheric flexure, and eustatic sea level change. Eustatic sea level changes due to climate changes in the

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Pacific Neogene will be discussed further below. N unn (1986) suggested that global changes in geoidal eustasy may explain previously anomalously distributed emergent reefs, but as changes in the geoid are extremely difficult to model (Momer 1976, Woodroffe 1988) and are unknown for the geologic past. "Local" tectonic uplift is particularly important along southwestern Pacific collisional coasts, such as (from south to north) along the coast of New Zealand (Schofield 1960), New Guinea (Chappell 1974, Aharon and Chappell 1986), Christmas Island (Woodroffe 1989), Tonga (Woodroffe 1988), the Philippines (Hashimoto 1981, Hashimoto et al 1984), the Marianas (Curray et al 1970), and Okinawa (Kawana and Pirazzoli 1985). Elastic deformation of the ocean floor beneath the load of water (Walcott 1972) and flexure of the lithosphere in response to volcanic loading (McNutt and Menard 1978, Spencer et al 1987) are possible regional scale compensatory movements and the latter seems to be the best explanation for several recently uplifted reefs in the Southeast Pacific (such as certain of the Cook Islands). Also recently suggested is uplift due to plate transport over aesthenospheric bumps or over plate-flex bulges on the ocean side of subduction zones (Scott and Rotondo 1983). Certain areas of the Pacific have been subject to differential subsidence. Schlanger and Premoli Silva (1981) and Schlanger et al (1987) suggested that the sea floor in the region of the N auru Basin had undergone several episodes of volcanism from the Aptian to the Campanian, resulting in lithospheric thinning and uplift. Thus Enewetak was formed in an area of sea floor that was anomalously shallow relative to its age (about 165 Ma), and the region has since subsided anomalously quickly relative to the predicted rates of sea floor subsidence of Parson and Sclater (1977). Based on their geophysical survey and correlation of the unconformity history of Bikini and Enewetak, Schlanger et al (1987) assume that this entire region has had a similar subsidence history since the Eocene.

Paleocenaography

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Changes in oceanographic patterns and eustatic sea level have an important influence on numerous aspects of a oceanic island and its biota, e.g. the existence, structure, and morphology of reefs and islands; productivity of lagoons; direction and perceived distance from source faunas; and temperature of the ambient surface waters, to name just a few. Likelihood of colonization, spedation, and local extinction, and thus local and regional diversity and disparity may all be influenced. Much of the major change in Indo-Pacific surface oceanography over the Neogene is tied to two key types of events, glaciation and progressive cutting of the Tethyan circum-equatorial current. Glaciation was developed in Antarctica during the Middle to Late Miocene (Kennett et al 1987) (see also Matthews and Poore 1980, Denton et al 1984, and Keigwin and Keller 1984, who claim Antarctic ice sheets were extensive by the Oligocene) and in the northern Hemispheric ice during the Plio-Pleistocene. The Tethyan current was cut first by the connection of Africa to Europe by 12 to 10 Ma (Adams 1989, Bonaduce 1983) and the movement of the Australian plate northward to block the Pacific from the Indian Ocean (Romine and Lombardi 1985, Savin et al 1985, Hutchinson 1988), and later by uplift of the Isthmus of Panama by about 2.5 to 2 Ma (Keller et al 1989). Many have interpreted oxygen isotope records to indicate that surface waters in the tropical Pacific warmed throughout the Miocene, while a west-east negative gradient decreased (Savin et al 1985). However, simultaneously, the latitudinal temperature gradient increased due to Antarctic glaciation, so that by the Late Miocene the higher latitude surface temperatures had become significantly colder (Savin et al 1975) and the tropical province had shrunk. The commencement of major Antarctic glaciation is believed by some to be indicated by a rapid change in oxygen-isotope values 15-13 Ma (Savin et al 1985, Woodruff 1985). The sea level record constructed by Vail et al (1977) and Haq et al (1987) from continental margins suggest in general a sea level rise of about 30m into the mid Miocene followed by a drop of 30m. Bourrouilh-Le Jan and Hottinger (1988) suggest that the Early Miocene sea level rise is responsible at least in part, through reef drowning, for many known cases of interrupted middle Miocene reef growth. Central

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Pacific reefs of the Middie Miocene were generally replaced by macroids of coralline red algae and encrusting acervulinid foraminifera. Sderactinian-dominated reef-growth resumed after a sea level drop in the Late Miocene. Based on benthic foraminiferal oxygen isotopes, the Latest Miocene and Earliest Pliocene may have been subject to gladal/interglacial cycles a t roughly one third the amplitude of Late Pleistocene signals (Hodell et al 1986). Hodell et al attribute this to fluctuations in the Antarctic ice sheet. A coincident Late Miocene 13C decrease (-6.2 Ma) has been attributed to changes in glacial inventory, deep-ocean circulation, a n d /o r productivity (Delaney and Boyie 1987 and references within). Based upon arguments using C d/C a ratios over the same interval, Delaney and Boyle argue that the change was likely in global inventory, due either to increased weathering of 13C-depleted rocks during regression an d /o r decreased deposition of organic carbon during sedimentation.

Shackleton et al (1984) showed that glacially-derived clastic sedim ents first occur in high latitude sediments by about 2.4-2.5 Ma, coincident with a large shift in delta 180 values (Shackleton and Opdyke 1976). This point is taken to represent the onset of large-scale Northern Hemispheric glaciation and rapid large-magnitude sea level oscillations. At 900,000 to 800,000 years ago the variability in am plitude of the signal changed and the dominant wavelength increased from 41,000 to 100,000 years (cf. Williams et al 1988); the change separates the Early and Late Pleistocene. Combined Late Pleistocene records of 180 in deep sea cores and uplifted Pacific coral reefs resulted in a detailed sea level curve for the past 240,000 years (Chappell and Shackleton 1986), particularly since the height of the last interglacial (approx 125,000 ybp). Pinter and Gardner (1989) furnished a continuous late Quaternary sea level curve for the past 140/300 y using a polynomial model which fits a series of fourth order equations to existing late Quaternary data. The history of the last glacial and interglacial stages have been studied intensively. Bloom et al (1974) and Aharon and Chappell (1986) showed, using coral reef terraces at Huon Peninsula, Papua New Guinea, that sea levels rose 5-6.5 m above present mean

13

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sea level during the last intergladal. Using oxygen isotopes from the giant clam Tridacna gigas, Aharon and Chappell (1986) suggest that temperature was similar to

present during the early interstadials of isotope stage 5 and 3 degrees below present during the later interstadials. Degladation and assodated changes in sea level have not been monotonic over the period since the last gladal maximum at 18,000 ybp (CLIMAP 1976), but have shown (1) a series of events in which gladal melting was more rapid (14-12,000, 10-9000, 86000 ybp) (Mix and Ruddiman 1985), (2) a brief interval of regladation (11-10,000 ybp) (Ruddiman and McIntyre 1981), and (3) a sea level maximum sometime in the mid to late Holocene. In the oceanic Padfic thL; maximum was up to 6m higher than present (Bloom et al 1974, Papua New Guinea; Streif 1978, Malacca Straits). Clark and Lingle (1979) and Isla (1989) suggest different regions of the Padfic have experienced different Holocene sea level variations; the region roughly south of the equator in the IndoPadfic has generally been predicted to reach a sea level maximum at 5000 ybp and gradually subside, while the region north of the equator shows a relatively constant rate of increase since 5000 ybp.

Geology of individual locations used in this study

Geology of Enewetak and fossil material studied

Enewetak Atoll consists of about 1400m of Eocene to modem carbonate sediments upon an early Tertiary or pre-Tertiary volcanic base (Emery et al 1954, Ladd 1973). One of the two deep cores (E-l) from the first drilling operations at Enewetak was found to have shallow-marine fossils throughout the entire interval; this is the core that confirmed Darwin's theory of atoll subsidence. The other core (F-l) passed through shallow marine sediments that could be correlated with E-l through the top 450m (1400 feet), below which it entered outer-slope sediments. Early studies suggested that important unconformities occur at the top of the Eocene, at the top of the Early Miocene, and above the top of the Late Miocene, which is now known to be

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Quaternary (Ladd 1973). Miocene sediments revealed pollen and spores of vegetation that are indicative of high-island forest conditions (Leopold 1970), as well as land snails that are typically found only on high limestone islands (Ladd 1958).

Three reference cores (KAR-1, OAR-2, OOR-17) approximately 350m long were drilled for the PEACE project reference cores in 1985 (Figure 1-3). The cores had 62, 76 and 52 % recovery, respectively. The upper 300m of these cores contain well-preserved fossils and were the subject of most of the fossil ostracode analyses in the chapters to follow. For the Defense Nuclear Agency, which funded the study, feet instead of meters were used in the large amount of literature generated by the program (see Wardlaw and Henry 1987, 1988). Thus for consistency with the bulk of this literature, from this point onward I will use feet to describe these particular cores. Ludwig et al (1988) converted feet to meters in their publications on Sr-isotope stratigraphy of the cores. Quinn (1989), in his study of carbon and oxygen isotopes, used the metric depth from the sediment-water interface of the lagoon, rather than from the top of the core. Henry et al (1986) presented the borehole core logs and lithologic descriptions, and Wardlaw and Henry (1986) described the physical stratigraphy of these cores. Wardlaw and Henry recognized three lithologic intervals which are separated by major unconformities in the middle Miocene and at the Early/Late Pliocene boundary. The three major packages can be split into 8 smaller ones based on 8 disconformities and discontinuities, many of which correspond to biostratigraphic zone boundaries (Cronin, in prep). These major disconformities represent subaerial exposure surfaces during sea level drops. Cronin et al (1986, in prep) divided the Middle Miocene to Holocene sequence into 12 zones based upon benthic microfossils (Figure 1-4). The combination of microfossils, disconformities, and lithostratigraphy provide a precise correlation among the Enewetak cores. Cronin (in prep) used benthic ostracode assemblages, which have been shown to be widespread in the tropical Pacific (Weissleader et al 1989, Cronin in press), to correlate the cores to other tropical Pacific locations where the ostracode fauna has been described. Bybell and Poore (in press) provided planktonic foraminiferal and nannofossil

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stratigraphy on which most of the absolute dates of this material are based.

Samples used from the PEACE cores

The samples used for assemblage and morphometric analyses are primarily from core OOR-17, which among the three reference cores has been sampled and picked for ostracodes m ost thoroughly and evenly. Samples from OOR-17 were studied at approximately 7-15 foot intervals over a depth of 1053 feet; the only important exception includes the interval between 300 and 500 feet which was largely recrystallized due to phreatic meteoric diagenesis (Quinn 1989) during a Late Pliocene sea-level drop.

Samples from core KAR-1, which is the next-best studied, were used

for part of the assemblage study. Specimens from KAR-1 and OAR-2 were used as supplementary material for morphometric analyses; specimens from the Early Pliocene interval of OAR-2 were especially useful since this section is less severely altered than in the other two cores (Quinn 1989). A list of samples used for the assemblage analyses from OOR-17 and KAR-1 can be found in Appendix 2. All bulk samples and microfaunal slides are stored in the Branch of Paleontology and Stratigraphy at the U.S.Geological Survey in Reston, Virginia.

Much effort has been concentrated on the Pliocene-Quatemaiy record of Enewetak, in the hope that it would provide a detailed record of the effect of sea level change on reef structure and in particular provide an uncomplicated record of sea level itself (Thurber et al 1965, Henry et al 1974, Tracey and Ladd 1974, Videtich and Tremba 1978, Szabo et al 1985, Quinn 1989). Much of the work on this interval at Enewetak owes its existence to two Department of Defense shallow drilling programs in the early 1970's, the PACE program (Henry et al 1974), which included two reef holes, and the EXPOE program (Couch et al 1975), which included 46 continuous cores drilled to about 50 to 90m. Couch et al (1975) described 6 stratigraphic intervals separated by 5 erosional

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unconformities from the EXPOE cores. Szabo et al (1985) reported radiocarbon and uranium series dates of Quaternary corals from 3 of the top 4 of these intervals, using shallow cores in the windward reef flat taken during the PACE program. These same intervals are recognizable by disconformities in the PEACE program reference cores and provide additional evidence for the absolute ages of sediments from the top 300' of the core (Wardlaw 1989, Bybell and Poore in press). Szabo et al suggest the dated intervals of reef growth represent 4 of the 5 interglacial high sea level stands recognized by Shackleton and Opdyke (1973): oxygen isotope stages 1 (Holocene), 5e (131,000+3030 ybp), probably 7 or possibly 9 (undated due to recrystallization); and 11, or possibly 9 or 13 (Ur series dated at 454,000+100,000 ybp). Quinn (1989) investigated the post-Miocene meteoric diagenesis of Enewetak, using delta ,80 and 13C signals, which provide distinct values during meteoric vadose and meteoric phreatic diagenesis. Quinn concluded that meteoric vadose diagenesis was not responsible for significant amounts of diagenesis beyond the subaerial exposure surface; however, meteoric phreatic diagenesis created mineralogic stabilization shortly after sediment deposition and during Late Pleistocene sea-level lowstands. The extensive recrystallization between 300' and 500' in the Enewetak cores is due primarily to the latter sort of diagenesis. According to Quinn temporally distinct phreatic lenses impart distinct geochemical and isotopic signatures on a regional (due to variations in host rock) and global (due to changes in ice volume and C 0 2) scale; Quinn thereby correlated several Enewetak cores based upon the isotope signatures near subaerial exposure surfaces. Quinn suggested that several biostratigraphic boundaries placed at lithologicallydefined disconformities are actually positions of paleo-phreatic lenses. In fact, there is only one discrepancy, and I will show using quantitative assemblage analyses (Chapter 2) that these biostratigraphic boundaries were well-placed at intervals of faunal turnover. Application of oxygen and carbon isotopes to Late Miocene-Early Pliocene sediments that have retained their aragonite chemistry may give a primary record of changes in lagoonal productivity and temperature. Matthews and Quinn have made

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the isotopic measurements, but are stil! working on the results (Quinn 1990, pers. comm.). Ludwig et al (1988) attempted to date the top 350m from core KAR-1 using 67S r/“Sr isotope ratios. Sr-isotopic ratios, which (in theory) show monotonic variation with time in sea water (Peterman et al 1970), provide the potential for an important new technique of dating carbonate sediments (DePaolo and Ingram 1985, DePaolo 1986, Hodell et al 1989). However, the results of the analyses of Ludwig et al (1988) differ from biostratigraphic and radiometric age estimates (Wardlaw 1989, Ross et al 1990). Hailey and Ludwig, in reply to Wardlaw, point out that Sr-isotope results are consistent with biostratigraphic estimates if one assumes a two million year uncertainty around biostratigraphic estimates. However, biostratigraphic (and radiometric) uncertainty is much less than two my for the top 100m of the cores, thus large deviations of Srisotopes from other data in this interval suggests deviations may be at least as large throughout the rem ainder of the cores. Strontium-isotope dating depends upon establishing an accurate and precise calibration curve; the Neogene calibration curve constructed by Hodell et al (1989) from 5 DSDP sites (three in the South Pacific and one each in the South Atlantic and Caribbean) show a great deal of scatter (Poore and Dowsett 1989). Kaufman et al (1989) and Poore and Dowsett (1989) provide short reviews of potential problems with Sr-isotope stratigraphy. Problems with the method have not been confined to the Enewetak cores, and may reflect inaccuracies in dating of the original calibration curves, in our understanding of the complexity of the curve, or in our understanding of organismic fractionation of strontium. At Enewetak it seems unlikely that the problem lies in "complex diagenesis" of sediments, as suggested by W ardlaw (Ludwig et al 1988). While the age inferred from the strontium ratios may be incorrect, the pattern of the curve, with several abrupt transitions, may be revealing with respect to periods of non-deposition (Ludwig and Hailey 1988).

Fossil material from outside Enewetak

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I washed and picked about a dozen samples of core or cuttings from Bikini core 2A (Emeiy et al 1954). Most of the material available was well lithified, and apparently much of the finer unconsolidated sediment from the cores was lost during drilling (Tracey 1989, pers. comm.). In this material I observed only a few samples with many ostracodes. In 1986 I collected from a cross section on the Calcar-Barilli road along the east central side of Cebu, Philippines, accompanied by people from the Philippine Bureau of Mines. Samples used here are from the Birili Formation, assigned by Maac (1986) from planktic foraminiferal data to the Late Miocene to Early Pliocene, and the Calcar Formation, which she assigned to Late Pliocene to Pleistocene. The Calcar rests unconformably upon the Birili. Muller et al (1980) came to similar conclusions; other interpretations exist (e.g. Hashimoto and Balce (1977), BMG (1981)). The samples of the Barili Formation from which I picked ostracode assemblages are from the Guadalupe Limestone, a shallowly deposited limestone member at the base of the formation. This member is mostly hard-bedded massive coralline limestone, but between these massive beds are thin (l-10cm) beds of chalky friable limestone. These friable units provide rich microfaunas, which appear to contain both shallow and shelfdepth assemblages. The samples from the Calcar Formation are from a very shallow limestone unit that has abundant quantities of the larger foraminifer Amphistegina and few planktic foraminifers; thus the age constraints on this unit appear to be much weaker. In 1986 the Australian Bureau of Mineral Resources undertook the North East Australia Project to core and dredge parts of the edge and slope of the Great Barrier Reef. I was given several samples from each of four cores taken through shallow water calcareous sediments. These cores drilled through Halimeda bioherms and probably contain Holocene or latest Pleistocene sediments. The calcareous faunas are species-rich and specimens are well-preserved. From the Bureau of Mines in Fiji I was able to sample several cores taken for hydrologic purposes that contain rich fossiliferous Neogene sediments. Core DDH(W) 78/2 consists of 200 feet of sediment from the Meighunyah beds on the western side of

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Viti Levu, the main island in Fiji (Rodda 1978). I also sampled from the Nakasi beds, 122m of largely shallow marine sediments taken near northeast of Suva, Viti Levu. The Nakasi beds overly the Early Pliocene Suva marl of the Meighunyah beds (Rodda 1981). In Taiwan I collected shallow marine mudstones and thinly-bedded carbonates in Hengchun Peninsula, southern Taiwan. The Plio-Pleistocene Maanshan Mudstone unconformably underlies the Pleistocene Hengchun Formation, which conformably underlies the Pleistocene limestones of the Ssoukou Formation (Hu 1984, Cheng and Huang 1975). Ostracode faunas from these formations were previously described by Hu (Mannshan Mudstone: Hu 1981a, 1982b; Hengchun Limestone: H u 1981a, 1982b; and Ssoukou Formation: H u 1984).

ECOLOGIC CONTEXT: LOCAL ECOLOGY AND GEOMORPHOLOGY OF MICRONESIAN LAGOONS

Island Physiography

Size and shape

Habitat area and heterogeneity are among the strongest ecological determinants of assemblage composition. Despite the supposed homogeneity in environment and fauna among the Micronesian islands, the land masses and accompanying lagoons vary in shape and size. Every major sort of oceanic island can be found in Micronesia, except for an island on a spreading ridge. The Marianas are all high islands. They vary in size from 0.9 to 541 sq km and are variously capped by limestone, volcanics, or both (such as at Guam). The resulting topography, drainage patterns, soils, and vegetation are radically affected by these variations in underlying rock, and so too are the marine reefs and faunas that develop around the edge of the island. Most of the islands are several hundred meters in elevation, reaching a maximum of 965m at Agrihan. Guam is the only island of this chain whose ostracodes were studied here.

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The Caroline Islands are especially interesting because they contain what might be considered the model Darwinian sequence of reefs, from the geologically recently erupted easternmost island Kusaie, to the half-atoll Pohnpei with its one central volcanic peak, to Truk with its huge lagoon and deeply dissected volcanic center islands, to full atolls in the west. In the far west are the lush and varied islands of Palau (of Jellyfish Lake fame). I used samples primarily from Truk, Pohnpei, and Pingelap. Pingelap is a tiny atoll-like structure in the east, between Pohnpei and Kusaie. The Marshall Islands are entirely atolls and vary in size from 5 sq km to the world's largest atoll Kwajalein (2300 sq km) (Emery et al 1954). The lagoons of m ost of the atolls discussed in this study (aside from Kwajalein), Enewetak, Bikini, Rongelap, Jaluit, Amo, and Majuro, are roughly the same size (500-1000 sq km); the land masses are only 5-10 sq. km. These would seem to be relatively homogenous ecologically in comparison to lagoons from other parts of Micronesia. The atoll islands are no higher than 4-6m in elevation (the highest point in Majuro is a 6m bridge connecting two islands), and the width of dry land is no wider than 200-400 meters between seaward and lagoonward beaches. The land area of Onotoa Atoll is greater than that of any of the Marshall Island atolls (13.5 sq km), but the lagoon is much smaller (425/>250/>180/>150, and b>c>d and e>f>g>h>i>.

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JACCARD COEFFICIENTS

Holocene vs. m o d em Enewetak

30 Intralagoonal Pleistocene vs. m odern Enewetak

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60 M iddle Miocene to Late Miocene vs. Late Pliocene to m odern Enewetak

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Figure 2-9. Jaccard coefficients between each sample and Quaternary samples from Enewetak. (A) Maximum values, (B) Average values. L: OOR-17 903'-800/; B: 776'-535', T: 299'-137', M: m odem Marshall Islands (excluding Guam), C: Caroline Islands, O: Onotoa, G: Guam. There is no trend within any of the three fossil groups toward m odem specimens, but there is a strong trend between them, especially over the gap before the appearance of modem-type assemblages in the Late Pliocene.

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Figure 2-9 cont'd. (C-E) Jaccard coefficients between one stratigraphic interval (bound by disconformities) and the one stratigraphically above it. (C) Jaccard coefficients between each sample from the Late Pliocene (OOR-17 299-137') and the group of four samples from the Pleistocene (OOR-17 lKy-70'), (D) Jaccard coefficients between each sample from the Pleistocene and and the group of samples from the Holocene (OOR-17 60'-14'), (E) Jaccard coefficients between each sample from the Holocene and the group of 4 m odern samples from Enewetak lagoon. (F) is the maximum and average Jaccard coefficient between successive stratigraphic group and the m odern samples from Enewetak lagoon (OOR-17 299'-137', 125', 110-70',60'-14', and intralagoonal). There is no trend in species composition within intervals toward the next, but there is a net trend from Late Pliocene to modem samples, primarily due to a shift at the Pleistocene/Holocene boundary.

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INTRA-ATOLL

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Figure 2-10. Species range chart. (A) Last appearance datums in Enewetak core OOR-17. (B) First appearance datums. These appearances and disappearance are local. No account was m ade for extinctions of taxa that later recolonized.

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OSTRACODE RANGE CHARTS ENEWETAK CORE O O R

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Figure 2-11. Species richness in m odem Micronesian lagoons. Lagoons with large islands in the center are the most diverse, Guam and Pohnpei. The least diverse were the true atolls of the Marshall Islands.

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MICRONESIA ESTIMATED LAGOONAL

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