Applications of Calcareous Nannofossils and Stable Isotopes To

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2007 Applications of Calcareous Nannofossils and Stable Isotopes to Cenozoic Paleoceanography: Examples from the Eastern Equatorial Pacific, Western Equatorial Atlantic and Southern Indian Oceans Shijun Jiang

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THE FLORIDA STATE UNIVERSITY

COLLEGE OF ARTS AND SCIENCES

APPLICATIONS OF CALCAREOUS NANNOFOSSILS AND STABLE ISOTOPES TO CENOZOIC PALEOCEANOGRAPHY: EXAMPLES FROM THE EASTERN EQUATORIAL PACIFIC, WESTERN EQUATORIAL ATLANTIC AND SOUTHERN INDIAN OCEANS

SHIJUN JIANG

A dissertation submitted to the Department of Geological Sciences in partial fulfillment of the requirements for the degree of Doctor of Philosophy

Degree Awarded: Fall Semester, 2007

The members of the Committee approve the Dissertation of Shijun Jiang defended on July 13, 2007.

______Sherwood W. Wise, Jr. Professor Directing Dissertation

______Richard L. Iverson Outside Committee Member

______Anthony J. Arnold Committee Member

______Joseph F. Donoghue Committee Member

______Yang Wang Committee Member

The Office of Graduate Studies has verified and approved the above named committee members.

To Shuiqing and Jenny

iii

ACKNOWLEDGEMENTS

First of all, I would like to thank my mentor Dr. Sherwood W. Wise, Jr., who constantly encouraged and supported me with his enthusiasm, reliance, guidance and, most of all, patience throughout my Ph.D adventure. He also opened a door into my knowledge of the paleo world, generously shared his time and his wealth of knowledge, patiently guided me through a western educational system totally different from my own background, and has successfully fostered my interest and enthusiasm in teaching. He deserves all my words of praise and appreciation.

I would also like to thank Drs. Anthony Arnold, Joseph Donoghue, Richard Iverson, and Yang Wang for their enthusiastic service on my committee and generously sharing their expertises and professional insights with me. Their support and insightful comments were invaluable at each stage of the dissertation process, and their contributions extend well beyond my dissertation. This long journey could not be so smooth and pleasant without their time and efforts.

Funding for research related to Ocean Drilling Program (ODP) Leg 207 was provided by a U.S. Science Support Program grant (Task Order F001790) to SJ and SWW, and general laboratory support provided by JOI USSAC (for ODP Legs 183 and 207 [Task Order T308A33]) and NSF-OPP grants 0126218 and 0230469 to SWW. This research used samples and/or data provided by the ODP. The ODP is sponsored by the U.S. National Science Foundation (NSF) and participating countries under management of Joint Oceanographic Institutions (JOI), Inc.

There are many other people and organizations to thank when you have gotten this far, especially during a more than five-year-long Ph.D adventure in a foreign country with very different systems and cultures from my own. They are as follows.

iv Much appreciation goes to the FSU Department of Geological Sciences for the financial support; to Nicole Bowens, Tami Karl, Mary Gilmore, and Ted Zateslo, thanks for taking care of those everyday but indispensable businesses during the process.

Special thanks go to the FSU International Center, specifically Kristen Hagen for her assistance and outstanding work in the clearance of visa issues that I seem to always have, both in my home country China and Republic of Panama.

Thank you to Velkis Candanedo, Carlos Langoni, Jeremy Brown (FSU Panama campus), Jack Baldauf (TAMU&ODP), Larry Obee (ODP) for their help when I got stuck with an expired visa in Panama. In particular, I would like to thank Velkis for her kind hospitability, accommodation, and fun tour around Panama City.

A special thank you extends to my ODP Leg 206 colleagues and the ODP 207 Shipboard Scientific Party for their relevant contribution to this research. To Gary Acton and Tariq Ayyub, thanks for your friendship; to other members of the Leg 206 Scientific Party, ODP Shipboard Personnel and Technical Representatives, and Transocean Drilling Crew, thanks for making our experience fun and interesting; to Richard Norris (Scripps Institution of Oceanography), thanks for providing samples for my PETM research from the U-channel through the archive half of the core and use of his unpublished bulk carbonate isotope data on the Leg 207 materials.

Invaluable tools brought much convenience and ease to my research. A big thank you to Mitch Covington (BugWare Inc.) for providing use of the patented nannofossil-counting program (BugWin 2K); to Dr. William Parker for the wonderful statistical programs; to FSU Nannofossil (Geology), Stable Isotope (Maglab) Laboratories (FSU Stable Isotope Lab has been supported by NSF grants EAR-0517806 and EAR-0236357), and Biological Science Imaging Resource (Biology), especially to Kristeen Roessig, Bryan Ladner, Audra Stant, Denise Kulhanek, Stacie Blair for help and discussion on sample preparation and exchange of ideas; to Yang Wang, Yingfeng Xu, and Dana Biasatti for facilitating the stable isotope analyses; and to Kim Riddle and Tom Fellers for assistance with the SEM work.

v The constructive and insightful comments and criticisms from the following reviewers greatly improved the manuscripts submitted for publication: Drs. Gary Acton, José-Abel Flores, and Koji Kameo for Chapter 1; Drs Isabella Raffi, Jeremy Young, and Jackie Lees for Chapter 2; Drs. Gerald J. Dickens, Mitchell W. Lyle, Larry C. Peterson, Gary Acton, and an anonymous reviewer for Chapter 3; two anonymous reviewers for Chapter 4; Drs Paul Bown and Jeremy Young for Chapter 5; Drs. Alan Cooper, Timothy Bralower, and John Barron for Chapter 6. Cinthia Wise and Nicolas Thibault assisted with the French translation for Chapter 4. James Arney provided a preliminary delineation of the PETM interval in Chapter 6.

Thanks to my friends here at Tallahassee, Hailin Deng and family, Chunfu Zhang and family, Xiaolong Hu and family, Xin Li, Yingfeng Xu, Shichun Huang, Xiaowei Jiang, Xiaojun Yang and Xiaode Deng, Xiaohang Zhang, the Furners (Chris and Zhan), Yi Cheng, Chunlin Huang, Zhiliang Wang, and my other volleyball fellows. I extend my gratitude for your friendship, sportship, and killing time together (fishing and poker), which together makes the stay in Tallahassee fun and enjoyable.

Last, but not least, I would like to thank my direct and extended family (mom, father-in-law, brother and sister-in-law, sisters and brothers-in-law, nephews, and nieces), who even though have no clue of what I am doing, for their unending love and positive support on my choices and my steps of life for the past years. Thank you to my wife, Shuiqing, for the years of company in a foreign country, and your love and encouragement along the way; to my daughter, Jenny, for her sweet smiles that help me see a bright future in lows of life and for her patience and tolerance to allow me to grow as a great father; to my father and grandfather, who unfortunately cannot see this.

TABLE OF CONTENTS

LIST OF TABLES...... x LIST OF FIGURES ...... xi LIST OF PLATES ...... xiii ABSTRACT...... xiv

INTRODUCTION ...... 1 Cenozoic Climate Evolution...... 1 Drill Sites, Objectives and Problems ...... 3 ODP Site 1256, Eastern Equatorial Pacific...... 3 ODP Site 1259, Demerara Rise, Western Equatorial Atlantic...... 5 ODP Site 1135, Kerguelen Plateau, Southern Indian Ocean...... 6 Paleoceanographic Approaches in Current Research ...... 7 Principal Results, Chapter Arrangement and Publication Status...... 10

1. UPPER CENOZOIC CALCAREOUS NANNOFOSSIL BIOSTRATIGRAPHY AND INFERRED SEDIMENTATION, ODP LEG 206, EAST PACIFIC RISE ..16 Abstract...... 16 Introduction...... 17 Materials and Methods...... 18 Biostratigraphy Results...... 19 Quaternary (0-1.77 Ma) ...... 20 Pliocene (1.77-5.32 Ma) ...... 21 Miocene (5.32 -23.80 Ma) ...... 23 Discussion...... 25 Sedimentation History ...... 25 Middle/Late Miocene Carbonate Crash...... 27 Downslope Transport of Sediments...... 28 Conclusions...... 30 Acknowledgements...... 30

vii 2. TAXONOMIC NOTE: DISCOASTER STELLULUS GARTNER, 1967, EMENDED ...... 62 Introduction...... 62 Materials and Methods...... 62 Systematic Palaeontology ...... 63 Acknowledgements...... 65

3. CAUSE OF THE MIDDLE/LATE MIOCENE CARBONATE CRASH: DISSOLUTION OR LOW PRODUCTIVITY?...... 67 Abstract...... 67 Introduction...... 68 Background...... 69 Study Area ...... 69 Samples, Methods, and Age Constraints ...... 70 Results: Variations in Bulk Isotopes and Carbonate Accumulation Rates...... 71 Discussion...... 72 Bulk-carbonate Isotopes...... 72 Carbonate Diagenesis...... 73 Surface-water Carbonate-production Proxies...... 74 Carbonate Crash...... 76 Timing of Carbonate Crash ...... 76 Carbonate Accumulation in Open Oceans ...... 76 Previous Models for the Carbonate Crash ...... 77 Proposed Model for the Carbonate Crash...... 79 Conclusions...... 82 Acknowledgements...... 83

4. SURFACE-WATER CHEMISTRY AND FERTILITY VARIATIONS IN THE TROPICAL ATLANTIC ACROSS THE PALEOCENE/EOCENE THERMAL MAXIMUM AS EVIDENCED BY CALCAREOUS NANNOPLANKTON FROM ODP LEG 207, HOLE 1259B...... 89 Abstract...... 89 Introduction...... 90 Materials and Methods...... 92 Results: Nannofossil Turnover across the PETM...... 94 Discussion...... 96 Environmental Impacts on Coccolith Formation...... 96 Correspondence Analysis Factors and Their Paleoenvironmental Indications...... 98 PETM Surface-water Environmental Variations...... 101 Response of Nannoplankton to the PETM...... 102 Conclusions...... 104 Acknowledgements...... 104

viii 5. TAXONOMIC NOTE: A NEW COCCOLITHUS SPECIES THRIVING DURING THE PALEOCENE/EOCENE THERMAL MAXIMUM...... 121 Introduction...... 121 Materials and Methods...... 126 Systematic Palaeontology ...... 126 Acknowledgements...... 126

6. ABRUPT TURNOVER IN CALCAREOUS-NANNOPLANKTON ASSEMBLAGES ACROSS THE PALEOCENE/EOCENE THERMAL MAXIMUM: IMPLICATIONS FOR SURFACE-WATER OLIGOTROPHY OVER THE KERGUELEN PLATEAU, SOUTHERN INDIAN OCEAN...... 127 Abstract...... 127 Introduction...... 128 Materials and Methods...... 129 Results: Variations in Nannofossil Assemblages ...... 130 Discussion...... 130 Stable isotope pattern...... 130 Ecological implications of assemblage variations ...... 131 Nannoplankton productivity changes across the PETM in the Southern Ocean ...... 132 Adaptive response of nannoplankton to the PETM ...... 132 Summary...... 133 Acknowledgements...... 133

APPENDIX A ...... 138

APPENDIX B ...... 142

REFERENCES...... 145

BIOGRAPHICAL SKETCH ...... 170

LIST OF TABLES

Table 1-1 Calcareous nannofossil distribution chart in Hole 1256B ……………………… 37

Table 1-2 Age estimate, depth, and zonation of calcareous nannofossil datum levels used in this study …………………………………………………………………….. 61

Table 4-1 Calcareous nannofossil distribution chart and counts ………………………… 111

Table 4-2 List of calcareous nannofossil taxa identified ………………………………… 114

Table 4-3 Results of correspondence analysis of relative abundance data ………………. 115

Table 6-1 Calcareous nannofossil counts in ODP Hole 1135A ………………………….. 137

LIST OF FIGURES

Figure 0-1 Cenozoic events in climate, tectonics, and biota vs. δ18O and δ13O in benthic foraminiferal calcite ……………………………………………………………..14

Figure 0-2 The studied ODP drilling sites superimposed on a plate tectonic map ………….15

Figure 1-1 Backtracked path for Leg 206 Site 1256 and locations of Leg 138 sites ………..32

Figure 1-2A Linear sedimentation rates as constrained by the biostratigraphy and magnetostratigraphy, and a best-fit linear sedimentation rate model …………... 33

Figure 1-2B Calcareous nannofossil datum age-depth plots for eastern equatorial Pacific drill holes ……………………………………………………………………….. 34

Figure 1-3 Relationship of mass accumulation rates between bulk sediments and the main sedimentary components ………………………………………………………...35

Figure 1-4 Correlation of nannofossil datums in EEP drill holes that reflects variations in sedimentation rates over time …………………………………………………... 36

Figure 3-1 Locations of Site 1256 and other ODP/DSDP sites with major world surface currents …………………………………………………………………………..84

Figure 3-2 (A) Variations of carbonate mass accumulation rates and stable oxygen and carbon isotopes in the studied interval, and (B) their correlation ……………….85

Figure 3-3 Relationship of mass accumulation rates between bulk sediments and the main sedimentary components ……………………………………………………..….86

xi Figure 3-4 CaCO3 and volcanic ash mass accumulation rates (MARs) from ODP Leg 165 ………………………………………………………………………………..…..87

Figure 3-5 δ13C of the modern western Atlantic, eastern Pacific, and circumpolar Antarctic ……………………………………………………………………..…..88

Figure 4-1 Map showing the location of ODP Site 1259 superimposed on a 55.5 Ma plate reconstruction …………………………………………………………………..106

Figure 4-2 Percentage abundance of selected calcareous nannofossil taxa and bulk carbonate δ13C ………………………………………………………………….107

Figure 4-3 Close-up photograph of the P/E boundary recovered in ODP Hole 1259B …... 108

Figure 4-4 Variations of nannofossil assemblage preservation and diversity across the PETM from ODP Site 1259 ……………………………………………………109

Figure 4-5 Results of the correspondence analysis ………………………………………...110

Figure 6-1 ODP Site 1135 and other reference sites on an ODSN 55-Ma plate reconstruction …………………………………………………………………..135

Figure 6-2 Percentage abundances of nannofossil groupings and bulk foraminiferal δ13C and δ18O across the PETM from ODP Site 1135 ………………………………136

xii

LIST OF PLATES

Plate 2-1 Paratype specimens of Discoaster stellulus Gartner, 1967, emend. Jiang & Wise, 2006 ……………………………………………………………………… 66

Plate 4-1 Calcareous nannofossils across the PETM from ODP Hole 1259B …………...116

Plate 4-2 Calcareous nannofossils across the PETM from ODP Hole 1259B …………...118

Plate 5-1 Coccolithus bownii Jiang and Wise, 2007 …………………………………….. 126

xiii

ABSTRACT

This dissertation is a collection of five calcareous nannofossil and one stable isotope studies on materials from Ocean Drilling Program (ODP) Legs 183 (Site 1135), 206 (Site 1256), and 207 (Site 1259) that target two important paleoceanographic events: 1) the middle/late Miocene carbonate crash, and 2) the Paleocene/Eocene Thermal Maximum (PETM). Site 1256 nannofossil biostratigraphy in Chapter 1 refined the author’s shore-based shipboard Quaternary-middle-Miocene nannofossil biostratigraphy with 16 zones/combined zones recognized based on 28 nannofossil datums. This chapter provides a chronologic framework for the age calibration of the first occurrence (7.18 Ma) and last occurrence (6.32 Ma) of Reticulofenestra rotaria, calculation of linear sedimentation rates, age determination of basalt basement (~14.5 Ma), and the recognition of the “carbonate crash” paleoceanographic event at the middle/late Miocene boundary. Reworked nannofossils and lithologic changes also allow a reading of a three-episode redepositional history (4.7, 8.3, and 10.7 Ma, respectively) in the eastern Pacific. The detailed examination of the Site 1256 material also yielded well-preserved Discoaster stellulus, for which only the distal view had been depicted in the original description. In Chapter 2, a redescription and re-illustration of both sides of this asterolith is provided. This should prevent misidentification of specimens in proximal view, thereby raising its potential application for middle-late Miocene biostratigraphy. Based on the above age model, in Chapter 3 stable oxygen and carbon isotopes were used for the first time to explore the late/middle Miocene “carbonate crash”. This carbonate transition is a widespread (eastern and central equatorial Pacific, Indian, South Atlantic, and the Caribbean), sharp decrease in carbonate mass-accumulation rates, which has previously been considered only 2 13 a dissolution event. The positive correlation (R = 0.75) between δ C and CaCO3 mass accumulation rates during 5-14 Ma at ODP Site 1256 clearly demonstrates that carbonate accumulation is mainly biologically controlled. The coincidence of the carbonate crash with

xiv negative excursions in δ13C and δ18O values suggests a causative mechanism related to surface- water productivity, as a result of surface-water warming and reduced upwelling. Based on these observations, one could speculate that the major middle/late Miocene sea-level drop may have caused the complete closure of the Indonesian Seaway, resulting in a piling-up of surface warm water in the west Pacific. The eastward spread of this nutrient-poor water then would have warmed sea-surface temperatures and reduced upwelling in the central and eastern Pacific, thereby creating a prolonged “El Nino” scenario and reducing biological productivity of phytoplankton. The reduction in carbonate supply to the deep waters consequently caused a rapid shoaling of the carbonate compensation depth, thereby triggering the carbonate crash. The PETM was a catastrophic, rapid greenhouse-forced global warming event ~55 m.y. ago that triggered an abrupt turnover in ocean chemistry and circulation as well as biota. Chapter 4 represents a quantitative study of the response of nannoplankton to the PETM at Demerara Rise, equatorial Atlantic (Site 1259). Toweius, Fasciculithus, and Chiasmolithus sharply decrease at the onset of the PETM, whereas Chiasmolithus, Markalius cf. M. apertus, and Neochiasmolithus thrive immediately after the event, which also signals the successive first appearances of Discoaster araneus, Rhomboaster, and Tribrachiatus. Two main environmental factors were extracted by correspondence analysis of relative abundance data. The time series of the two factors shows that during the PETM, 1) environmental stress (most likely from changes in seawater pH) increased and may well have also induced the evolution of ephemeral nannofossil “excursion taxa”; and 2) surface-water productivity increased at this site presumably due to higher runoff from continental areas. The local phytoplankton opportunist, Markalius cf. M. apertus, is described as a new species in Chapter 5, which will be published under the name Coccolithus bownii. Results presented in Chapter 6 from Site 1135 on the Kerguelen Plateau, Southern Ocean suggest that nannoplankton responded differently to the PETM at southern high latitudes. The onset of the carbon isotope excursion occurs within an 18-cm interval (instead of 1-2 cm as observed in most deep-sea sections) before the peak is reached, displaying a linear mixing curve. This indicates that the release of light carbon was a gradual, single injection, instead of multiple pulses as suggested in previous work, and that this sequence is highly expanded as a result of high sedimentation rates at this relatively shallow oceanic site. This is evidenced by the high

xv numbers of dissolution-susceptible holococcoliths (Zygrhablithus bijugatus) preserved throughout the sequence. Although r- and K-selected specialists exponentially increase in abundance at the onset, Chiasmolithus abruptly drops but then rapidly recovers, whereas Discoaster and Fasciculithus show opposite trends, indicating that in high latitudes, surface-water oligotrophy prevailed at the carbon isotope excursion (CIE) onset but mesotrophic conditions dominated the CIE recovery. These observations confirm previous results from ODP Site 690 on Maud Rise. The intensive dissolution of susceptible holococcoliths and the poor preservation of the assemblages are believed to have been caused by the effects of corrosion caused by the methane release. The different responses of nannoplankton to the PETM and the contrasts evident in previous work from the open ocean vs the continental margins further demonstrate that the response to the PETM can be influenced by local differences in geologic setting and oceanographic conditions.

xvi

INTRODUCTION

Cenozoic Climate Evolution

The Cenozoic Era spans the last 65 m.y. during which most mammalian species evolved. During the Cenozoic, the Atlantic and Indian oceans continued their expansion at the expense of the Pacific Ocean, and the once virtually continuous equatorial seaway became progressively segmented by the closure of the Tethys (early to middle Cenozoic), blocking of the Indonesian Seaway (middle to late Cenozoic) due to northward migration of Australia, and the bridging of the Middle American Seaway (Pliocene). This process led to the diminution and final destruction of global equatorial circulation, which coincided, in general, with the development of the Antarctic Circumpolar Current (ACC) system via the opening of the Tasmanian Seaway and the Drake Passage. The ACC thermally isolated Antarctica by decoupling the warmer subtropical gyres from the colder subantarctic and Antarctic gyres, which, in turn, led to the development of cryospheric circulation at southern high latitudes. These events had substantial influence on the evolution of Cenozoic climates (Kennett, 1982). The evolution of Earth’s climate during the Cenozoic exhibits a long general cooling trend from a “greenhouse” state to the present “icehouse” state (Fig. 0-1). This climatic cooling, however, was not a smooth, gradual decrease of temperature, but consisted of a series of rather rapid, stepwise transitions punctured by a number of global climatic extremes (Fig. 0-1). Among these the most pronounced global carbon-cycle perturbation and warming in recent Earth history is associated with the Paleocene-Eocene Thermal Maximum (PETM) ~55.5 m.y. ago (e.g., Kennett and Stott, 1991). Most fingers of blame point to a skyrocketing amount of greenhouse gases in the atmosphere that is most parsimoniously explained by a massive dissociation of sedimentary methane hydrates stored along continental margins (e.g., Dickens et al., 1995, 1997). This event has been

1 widely accepted as one of the best examples of deep-time, rapid greenhouse-forced warming, and has remained one of the hottest topics of debate and study since its first discovery at Maud Rise, Weddell Sea, particularly in view of current worldwide concerns about global warming (see Chapters 4 and 6 herein). Thereafter, Paleogene climate cooled down and eventually operated under a fully glacial mode during the Neogene, which is believed to have been caused by the Himalayan uplift generated by the Indian subcontinent ramming into Asia (e.g., Raymo and Ruddiman, 1992). During the early Neogene, open circulation between equatorial Pacific and Atlantic oceans remained feasible via the Middle American Seaway. The Tethys Seaway, however, was increasingly restricted by the northward drift of Australia, and the Drake Passage opened further to accommodate deep-water circulation (e.g., Ciesielski and Wise, 1977; Barker, 2001). This amplified the Antarctic Circumpolar Current and thus intensified the thermal isolation of Antarctica. By 15 Ma, the permanent emplacement of the East Antarctic Ice Sheet had been established (Shackleton and Kennett, 1975; Savin et al., 1985), leading to 1) enhanced dissolution of carbonate, as indicated by an abrupt increase in the number of hiatuses in the equatorial Pacific (one of which is associated with the middle/late Miocene carbonate crash; see Chapter 3), and 2) increased fertility near the equator from intensified upwelling (van Andel et al., 1975a, b; Shackleton et al., 1997). During the Plio-Pleistocene, continued cooling climate resulted in northern Hemisphere glaciation, a series of “ice ages” with interspersed warm periods (e.g., Berggren and van Couvering, 1979; Shackleton et al., 1988; Maslin et al., 1998). The equatorial Pacific and Atlantic Oceans have long been known to be important repositories of information about Earth’s history. Sediments that have accumulated beneath the divergence-driven upwelling system provide the most continuous high- resolution archives and sensitive recorders of oceanographic change. These document a complex interplay of ocean chemistry, productivity, climate, ocean-atmospheric circulation, and plate tectonics (van Andel et al., 1975a, b; Shackleton et al., 1997). Because the sediments are shallow-buried, crop out on the seafloor in places, and contain spectacularly well preserved microfossils, they are ideal for studying Earth’s past oceanographic and climatic fluctuations.

2 Drill Sites, Objectives and Problems

The primary objective of this dissertation is to interpret data from calcareous nannofossil assemblages and stable oxygen/carbon isotopes to reconstruct the paleoceanographic conditions for Neogene and Paleogene materials from Ocean Drilling Program (ODP) Legs 183, 206, and 207 (Fig. 0-2). The writer participated on ODP Cruise 206 as a calcareous-nannofossil micropaleontologist. Detailed objectives for each expedition are as follows.

ODP Site 1256, Eastern Equatorial Pacific ODP Leg 206 was the first part of a two-leg campaign to sample a complete crustal section into the uppermost basement gabbros formed at a superfast spreading rate (Shipboard Scientific Party, 2003a). Site 1256 lies in 3635 m of water depth in the Guatemala Basin on Cocos plate crust formed about 15 Ma ago on the eastern flank of the EPR when this site experienced a super-fast spreading rate (Fig. 0-2; Wilson, 1996). This site lies at an equatorial latitude within the equatorial high-productivity zone and initially experienced very high sedimentation rates (>35 m/m.y.) (e.g., Shipboard Scientific Party, 2003a). During Leg 206 at Site 1256, three holes were drilled of which Hole 1256B cored a complete sedimentary sequence overlying oceanic basement. The sediments are subdivided into two principal lithologic units. Unit I (0-40.6 mbsf) is clay-rich with a few carbonate-rich intervals (becoming increasingly calcareous with depth), whereas Unit II (40.6-250.7 mbsf) is predominantly biogenic carbonate with some minor, more siliceous and diatom-rich intervals. Bioturbation is common throughout the whole sedimentary sequence. Calcareous nannofossil assemblages are generally abundant and moderately to well preserved at this site. Chronology is a key component of geologic studies. Biostratigraphy provides a useful tool for dating not only sedimentary marine sequences but also the underlying oceanic basalts. Dating the latter has relied heavily on the age of the sediments immediately above the crust because biostratigraphy has proven to be the most cost- and time-efficient dating method, whereas direct dating of oceanic basalts has seldom been

3 possible. Magnetostratigraphy has developed into a most important tool but because a magnetic reversal has no unique signature, it must be identified by other means such as micropaleontologic methods or stratigraphic correlation (e.g., by reference to a section on land that has radiometric tie-points). At Site 1256 in particular, biostratigraphy provides better constraints than magnetostratigraphy because, 1) the low paleolatitude of this site makes it nearly impossible to determine magnetic polarity from azimuthally unoriented core samples; 2) the nearly north-south spreading ridge orientation makes the magnetic inclination insensitive to structural tilting; 3) below 110 mbsf especially, the paleomagnetic data are very noisy due to a combination of weak intensities, drilling overprints, drilling deformation, and possibly some level of reduction diagenesis; and 4) there is an absence of polarity reversals in the upper Pleistocene. The routine shipboard biostratigraphy was derived from core-catcher samples with a ~9-m sample spacing, which provided low biostratigraphic resolution and therefore compromised the dating of the underlying basement, hence there was a strong need for a detailed nannofossil biostratigraphy based on shorter-spacing samples. A revised biostratigraphy based on a 75-cm spacing is given in Chapter 1. Excellent preservation of specimens of Discoaster stellulus in this material allowed a redescription and emendation of this species based on illustrations from the scanning electron microscope (SEM) in Chapter 2. A sharp decrease in sedimentation rates and carbonate concentration at the middle/late Miocene transition was noted on the biostratigraphic age-depth plot. Although the name "Carbonate crash" was applied to this particular paleoceanographic event a decade ago (Farrell et al., 1995; Lyle et al., 1995) and the phenomenon was known before that, its cause has been a subject of speculation. This event has also been reported in other areas (e.g., Caribbean, south Atlantic, and Indian Ocean) far removed from the original study area (equatorial Pacific), and no previous attempt has been made to explore the causative mechanism(s) in light of stable isotopes despite the feasibility and availability of such data. As explained and developed in Chapter 3, however, the generally high sedimentation rates and sediment recovery at Site 1256 has allowed a high-resolution study and a new interpretation of this widespread event.

4 ODP Site 1259, Demerara Rise, Western Equatorial Atlantic One of the primary scientific objectives of ODP Leg 207 was to evaluate the biotic turnover across the PETM. During the expedition, five sites were cored on the northern margin of the Demerara Rise at water depths from 1900 m to 3192 m (Fig. 0-2), and the Paleocene/Eocene boundary (a dark green clay-rich bed in sharp contact with underlying chalk) was recovered at each site. This clay layer is provisionally interpreted as the record of the shoaling of the carbonate compensation depth (CCD) following the P/E boundary event and is expected to coincide with both the benthic foraminiferal extinction events and the carbon-isotopic excursion (Shipboard Scientific Party, 2004a). “Excursion taxa”, those exclusively confined to the PETM, have been identified among the planktonic foraminiferal assemblages (Norris, 2001, pers. comm.; See Kelly et al., 1996 for definitions). Surface-dwelling phytoplankton appear to have responded to the PETM environmental changes in fundamentally different ways from benthic organisms (e.g., 30%–50% of benthic foraminifers went to extinction), but their strategies remain poorly understood due to the limited number of high-resolution investigations using calcareous nannofossils (e.g., Monechi et al., 2000; Bralower, 2002; Gibbs et al., 2006) and geochemical proxies (e.g., Sr/Ca ratios and biogenic Barium accumulation rates [Stoll and Bains, 2003; Bains et al., 2000]). An oligotrophic condition was suggested to have prevailed in the southern high- latitude open ocean during the PETM (Bralower, 2002). However, it remains poorly known whether or not phytoplankton respond in same way in different nutrient settings, e.g., on continental shelves and margins where presumably increased riverine runoff would raise the supply of nutrients and trace elements that could potentially increase the productivity. Accurate evaluation of the phytoplankton response to the PETM requires not only multi-proxy investigations at same site (e.g., Site 690) to eliminate the limitation of an individual proxy, but also comparable investigations at localities with different nutrient regimes to achieve a balanced picture. The low-latitude location of Site 1259 has provided an ideal opportunity to look into the impact of different oceanographic settings on the nannoplankton response to the PETM, as well as the shelf-to-open-ocean nutrient

5 gradient. These subjects are explored in Chapter 4. The local opportunistic species thriving during the PETM, Markalius cf. M. apertus, is described as a new species (Coccolithus bownii) in Chapter 5.

ODP Site 1135, Kerguelen Plateau, Southern Indian Ocean The overall objective of Leg 183 was to evaluate the hypothesis that the Kerguelen Plateau and Broken Ridge Large Igneous Province formed as a result of volcanism associated with ascent of a plume head and subsequent volcanism resulting from the plume stem (Frey et al., 2003). For these studies chronology is crucial, either for dating the oldest sediment above basaltic basement or more direct, post-cruise 40Ar/39Ar dating of the lavas, to understand the magma flux and the involved physical and chemical processes. ODP Site 1135 lies in 1567 m of water on the Kerguelen Plateau (Fig. 0-2). A low- resolution, preliminary study by Arney (2002) indicated the possible presence of the PETM in ODP Core 183-1135A-25R-4, a carbonate-rich nannofossil and foraminifer ooze without the characterized boundary clay layer as seen in low latitudes. In chapter 6, the presence of the PETM is confirmed by bulk foraminiferal isotopes and nannofossil assemblages at this southern high-latitude site. Contrasts among studies from open ocean and continental margins/shelves suggest that local differences in geologic setting and oceanographic conditions influence the biotic response to the PETM (Bralower, 2002; Gibbs et al., 2006; Jiang and Wise, 2006). However, controversy concerning the phytoplankton’s response to the PETM has resulted from studies of the classic locality on Maud Rise in the Southern Ocean that use different productivity proxies (Bains et al., 2000; Bralower, 2002; Stoll and Bains, 2003). Geochemical proxies, e.g. Sr/Ca ratios and biogenic barium accumulation rates, can be complicated by diagenesis and their affiliation with methane hydrates (Dickens et al., 2002; Bralower et al., 2004), while micropaleontologic assemblages may be subject to selective preservation and dissolution. In Chapter 6, a high-resolution study of nannofossils and bulk stable isotopes helps unveil the actual impact of the PETM on surface-water dwellers at the high southern latitudes.

6 Paleoceanographic Approaches in Current Research

The ocean floor, except for certain very young areas at the mid-ocean ridges, is covered with sedimentary strata and sediments in different thickness, great at the continental margins and little in the abyssal plains. They record information about the history of the earth’s environment and life, most of which can only be decoded using indirect proxies. There are many proxies in paleoceanography to assess the past ocean’s environmental and biological conditions (e.g., temperature, salinity, seawater pH, surface and deep circulation, water-mass distribution, food chain, productivity, etc). Among these, microfossils and stable isotopes are commonly used. Only a few planktonic and benthic organisms in the ocean have hard parts and, therefore, a chance of fossilization as major contributors of material to deep-sea sediments. Among them, microorganisms such as foraminifers, nannoplankton, diatoms, radiolarians, and dinoflagellates are of fundamental importance because their 1) small size creates a better chance of recovery in cores, 2) large numerical abundance renders them suitable for the application of more rigorous quantitative methods of analysis, and 3) wide geographic distribution makes them indispensable for regional correlations and comparisons, and paleoceanographic reconstructions. Specifically, calcareous nannofossils have been employed in this research. Their living descendants, coccolithophorid algae or nannoplankton, live in the photic zone and respond to environmental parameters (salinity, temperature, nutrients, etc.). Their modern distribution patterns of various forms have been widely studied in order to relate that distribution to oceanographic conditions (e.g., McIntyre and Bé, 1967; Okada and Honjo, 1973; Ziveri, et al., 1995). Past biogeography of calcareous nannofossils has also been widely investigated in an effort to interpret paleoceanographic conditions (e.g., Haq and Lohmann, 1976; Haq, 1980; Wei and Wise, 1990; Chepstow-Lusty et al., 1992; Monechi et al, 2000). Statistical analyses of nannofossil assemblages, combined with sedimentological, geochemical and other information, have helped establish paleoecological parameters for extinct taxa (Wei and Wise, 1990; Bralower, 2002; Jiang and Wise, 2006).

7 Historically, the past 38 years have witnessed the development of extensive applications of micropaleontology to the study of paleoceanography. During this time, the Deep Sea Drilling Project (DSDP) and ODP initiatives drilled more than 1300 sites in every ocean basin in the world including the Arctic and brought up Cenozoic and Mesozoic sediment cores from the great majority of them. Most of the sediments were dated first by micropaleontologic methods. Many scientific highlights have been obtained relative to plate tectonics, climate change, and sea-level change. As the oceans occupy 72% of the surface of the globe, the results were bound to have had a profound effect on geologic thinking. Analyses of stable oxygen and carbon isotopes have played a pivotal role in paleoceanography since the pioneering efforts of Emiliani (1955), who, building on work of Urey (1947), McCrea (1950) and Epstein et al. (1953), interpreted the isotopic record from deep-sea cores as a proxy for a series of Pleistocene climate/temperature cycles. As variations in oxygen isotopes recorded by planktonic foraminifera show similar trends throughout the world oceans, oxygen isotope stratigraphy has become not only a global correlation tool, but an established dating tool. δ18O signals from deep-sea benthic foraminifers, which avoid the “noise” imposed by short-term temperature and salinity variations, reflects fluctuations predominately of global ice volume and secondarily of temperature (Shackleton and Opdyke, 1973). This discovery resulted in widespread use of oxygen isotopic records in paleoceanography. Turning attention to carbon isotopes, Shackleton (1977) showed the potential significance of δ13C in studying water mass movement and paleoproductivity, and believed there is a connection between climatically induced changes in the terrestrial biosphere with observed carbonate dissolution cycles and the flux of dissolved CO2 in the ocean. Inspection of composition fluctuations of oxygen and carbon isotopes in carbonate from the sediments, therefore, can provide not only age constraints, but also valuable information for paleoceanography. Stable isotopic applications in paleoceanography usually focus on oxygen and carbon isotopes, i.e., the 18O/16O and 13C/12C ratios. Although each pair of isotopes has identical chemical properties, basically, molecules consisting of light isotopes react more easily than those consisting of heavy isotopes, and cause either equilibrium fractionation or kinetic fractionation. The former is essentially temperature dependent, but the later is

8 much more complicated. During physical and/or chemical processes, the light isotopes are fractionated out; therefore, heavy isotopes become relatively enriched in the seawater, where marine organisms such as foraminifers and nannoplankton record such changes in their calcite skeletons. The seawater isotopic signature becomes fossilized when the organisms die and are deposited. Seawater oxygen isotopic changes are closedly linked with the hydrological cycle. Although the common light 16O and the rare heavy 18O have identical chemical properties, the light one exhibits higher vapor pressure. Thus evaporation causes 16O to be preferentially higher in water vapor and 18O to be relatively enriched in seawater, which renders variations in δ18O. These variations are useful in studies of late Cenozoic paleoceanography since large polar and mountain ice sheets lock up large amounts of freshwater and their volumes fluctuate in time. Carbon exists in two main reservoirs: organic matter and sedimentary carbonates. In the modern ocean, the distribution of 13C is mainly controlled by the biological pump.

CO2 is “drawn down” from the atmosphere to the ocean by diffusion, and subsequently this inorganic carbon is converted into organic matter by photosynthesis of phytoplankton in ocean’s photic zone. Marine phytoplankton produce organic matter with negative δ13C values (-20~-30 ‰; ~-50 ‰ for bacteria-generated biogenic methane; Trumbore and Druffel, 1995) relative to ambient water due to the strongly preferential uptake of 12C, which causes surface-water dissolved carbon relatively enriched with 13C. These signals can be recorded in coccoliths, foraminiferal tests, and other kinds of carbonate formed therein that fossilize when the biogenic carbonates are buried. Therefore, variations in δ13C values can reflect the surface-water productivity, invaluable information for understanding of ocean/atmosphere system (an important component of the global carbon cycle). Variations in δ13C of benthic foraminifers, on the other hand, provide information about processes associated with the global carbon cycle and deep circulation that might trigger or arise from climate changes.

9 Principal Results, Chapter Arrangement and Publication Status

The results of my dissertation are compiled in different chapters in paper format (following the journal in which the chapter was published or will be submitted) as follows:

Chapter 1: Upper Cenozoic Calcareous Nannofossil Biostratigraphy and Inferred Sedimentation, ODP Leg 206, East Pacific Rise. Principal results: This chapter is basically the necessary shore-based refinement of the shipboard nannofossil biostratigraphy. Sixteen zones/combined zones were recognized based on 28 nannofossil datums. The established biostratigraphy as well as magnetostratigraphy were used to calibrate the ages of the first occurrence (7.18 Ma) and last occurrence (6.32 Ma) of Reticulofenestra rotaria, to calculate linear sedimentation rates, and to derive the age of basalt basement (~14.5 Ma). Based on the pattern of variations in sedimentation rates, the “carbonate crash” was located at the middle/late Miocene boundary. Reworked nannofossils and lithologic changes were used to unravel redepositional history, and three episodes were recognized (4.7, 8.3, and 10.7 Ma, respectively), with one at 8.3 Ma regionally correlated. Authors: Shijun Jiang and Sherwood W. Wise Journal: Proceedings of the Ocean Drilling Program, Scientific Results, 206: 1–25. doi:10.2973/odp.proc.sr.206.008.2007 Status: Published on January 9, 2007.

Chapter 2: Taxonomic note: Discoaster stellulus Gartner, 1967, emended. Principal results: This chapter includes a redescription and re-illustration of Discoaster stellulus, a common middle-late Miocene nannofossil that was originally described in 1967 largely based on the distal view of the asterolith (the proximal view not depicted). The emended description is based on extensive examination of specimens under light microscopy and SEM. A mobile-mount technique was used to observe multiple views of same specimens. This redescription will help

10 prevent misidentification of specimens in proximal view, thereby raising its potential application in biostratigraphy. Authors: Shijun Jiang and Sherwood W. Wise Journal: Journal of Nannoplankton Research, 28(2): 81–83. Status: Published in September 2006.

Chapter 3: Cause of the middle/late Miocene carbonate crash: dissolution or low productivity? Principal results: Stable oxygen and carbon isotopes were used for first time to explore the late/middle Miocene “carbonate crash”, a widespread (eastern and central equatorial Pacific, Indian, south Atlantic, and the Caribbean), sharp decrease in carbonate mass-accumulation rates that has previously been considered only a 2 13 dissolution event. The positive correlation (R = 0.75) between δ C and CaCO3 mass accumulation rates from 5-14 Ma at ODP Site 1256 demonstrates that carbonate production is mainly biologically controlled. The coincidence of the carbonate crash with negative excursions in δ13C and δ18O values suggests a causative mechanism related to surface-water productivity, as a result of surface- water warming and reduced upwelling. Based on these observations, one could speculate that the major middle/late Miocene sea-level drop may have caused the complete closure of the Indonesian Seaway, resulting in a piling-up of surface warm water in the west Pacific. The eastward spread of this nutrient-poor water then would have warmed sea-surface temperatures and reduced upwelling in the central and eastern Pacific, thereby creating an “El Nino” scenario and reducing biological productivity of phytoplankton. The reduction in carbonate supply to the deep waters consequently caused a rapid shoaling of the carbonate compensation depth, thereby triggering the carbonate crash. Authors: Shijun Jiang, Sherwood W. Wise, and Yang Wang Journal: Proceedings of the Ocean Drilling Program, Scientific Results, 206: 1–24. doi:10.2973/odp.proc.sr.206.013.2007 Status: Published on April 30, 2007.

11 Chapter 4: Surface-water chemistry and fertility variations in the tropical Atlantic across the Paleocene/Eocene Thermal Maximum as evidenced by calcareous nannoplankton from ODP Leg 207, Hole 1259B. Principal results: This chapter consists of a quantitative study of nannofossil data using statistical approaches. Toweius, Fasciculithus, and Chiasmolithus sharply decrease at the onset of the PETM, whereas Chiasmolithus, Markalius cf. M. apertus, and Neochiasmolithus thrive immediately after the event, which also signals the successive first appearances of Discoaster araneus, Rhomboaster, and Tribrachiatus. Two main environmental factors were extracted by correspondence analysis of relative abundance data. The time series of the two factors shows that during the PETM, 1) environmental stress (most likely seawater pH) increased and may well have also induced the evolution of ephemeral nannofossil “excursion taxa”; and 2) surface-water productivity increased at this site presumably due to higher runoff from continental areas. Contrasts between the results of this study and previous work from open ocean and continental margins further demonstrate that the response to the PETM can be influenced by local differences in geologic setting and oceanographic conditions. Authors: Shijun Jiang and Sherwood W. Wise Journal: Revue de Micropaléontologie, 49: 227–244. doi:10.1016/j.revmic.2006.10.002 Status: Published in October 2006.

Chapter 5: Taxonomic note: A new Coccolithus species thriving during the Paleocene /Eocene Thermal Maximum. Principal results: This chapter describes a new species, Coccolithus bownii that thrived during the PETM. It is a nearly round Coccolithus with a large central opening, and has a herringbone birefringence pattern characteristic of the genus. A freshly- broken facture section was used to yield clay-free specimens under the SEM. Authors: Shijun Jiang and Sherwood W. Wise Journal: Journal of Nannoplankton Research. Status: Submitted.

12 Chapter 6: Abrupt turnover in calcareous-nannoplankton assemblages across the Paleocene/Eocene Thermal Maximum: implications for surface-water oligotrophy over the Kerguelen Plateau, Southern Indian Ocean. Principal results: ODP Core 183-1135A-25R-4 from the Kerguelen Plateau in the Southern Ocean represents only the second complete, expanded section of the PETM, as defined by a carbon isotope excursion. Although Chiasmolithus, Discoaster, and Fasciculithus exponentially increase in abundance at the onset, the former abruptly drops but then rapidly recovers, whereas the latter two taxa show opposite trends due to surface-water oligotrophy. These observations confirm previous results from ODP Site 690 on Maud Rise. The intensive dissolution of susceptible holococcoliths and the poor preservation of the assemblages are believed to have been caused by a rapid ocean acidification. Authors: Shijun Jiang and Sherwood W. Wise Journal: Proceedings Volume: Antarctica: A Keystone in a Changing World – Online Proceedings of the 10th ISAES (International Symposium on Antarctic Earth Sciences). Status: Published in July 2007.

Figure 0-1. Cenozoic events in climate, tectonics, and biota vs. δ18O and δ13O in benthic foraminiferal calcite (from Zachos et al., 2001). VPDB = Vienna Peedee belemnite.

Figure 0-2. The studied ODP drilling sites (solid red circles) superimposed on a plate tectonic map of the world (after Shipboard Scientific Party, 2003b).

CHAPTER 1

UPPER CENOZOIC CALCAREOUS NANNOFOSSIL BIOSTRATIGRAPHY AND INFERRED SEDIMENTATION, ODP LEG 206, EAST PACIFIC RISE

Abstract

Site 1256 of Ocean Drilling Program (ODP) Leg 206 to the Guatemala Basin on the eastern flank of the East Pacific Rise yielded a near-complete, middle Miocene through Quaternary carbonate-rich section that provides an opportunity to study low-latitude biostratigraphic and paleoceanographic events. The sedimentary sequence in Hole 1256B has been zoned using calcareous nannofossils according to the biostratigraphic schemes by Martini of 1971, Martini and Müller of 1986, and Okada and Bukry of 1980. The nannofossil assemblage is characteristic of the low-latitudes, with abundant Gephyrocapsa, Discoaster, and Sphenolithus, and, is in general moderately- to well-preserved depending on nannofossil abundance and the presence of diatoms. Age estimates for the first occurrence and last occurrence of Reticulofenestra rotaria were derived from biostratigraphy and magnetostratigraphy independently and assigned to 7.18 Ma and 6.32 Ma, respectively. Linear sedimentation rates, calculated using 28 nannofossil datums and age estimates, are high in the middle Miocene, decrease from the late Miocene to the Pliocene, then increase upsection. The abrupt drop in carbonate mass accumulation rates during early late Miocene is referred to as the “carbonate crash”. This pattern reflects (1) the long-trend decrease of productivity as the site moves away from the upwelling system at the equatorial divergence as well as (2) fluctuation in the chemistry of the bottom waters associated with production of the

16 North Atlantic Bottom Water and ventilation via the Panama Gateway. A basement age of 14.5 Ma was obtained by extrapolating the 39.1-m/m.y. rate in the middle Miocene to the basement at 250.7 meters below seafloor, and is consistent with the ~15-Ma age of the oceanic crust estimated from marine magnetic anomalies. Reworked nannofossils and lithologic changes were used to unravel postdepositional history, and three episodes were recognized, one of which in the latest Miocene can be widely correlated. Key words: calcareous nannofossils, biostratigraphy, sedimentation rates, redeposition, eastern equatorial Pacific

Introduction

Ocean Drilling Program (ODP) Leg 206 is the first of a two-leg campaign to sample a complete crustal section from the extrusive lavas, through the sheeted dike complex, and into the uppermost gabbros within in situ ocean crust, by drilling a deep hole into basement formed at a super-fast spreading rate (Shipboard Scientific Party, 2003c). Site 1256 (6°44.2’N, 91°56.1’W; Fig. 1-1) lies in 3635 m of water in the Guatemala Basin on Cocos-plate crust that formed ~15 m.y. ago on the eastern flank of the East Pacific Rise (EPR) when the site experienced super-fast spreading (Wilson, 1996). This site formed at an equatorial latitude (Fig. 1-1) within the equatorial high-productivity zone and initially endured very high sedimentation rates (>35 m/m.y.) (e.g., Shipboard Scientific Party, 2003c). Three holes were drilled at Site 1256 with an intention to retrieve sediments. Hole 1256A produced a single mudline core that recovered the uppermost 2.37 m of the sedimentary section. This interval was re-sampled using the advanced piston corer (APC) with a second mudline core in Hole 1256B, where coring continued to a depth of 250.70 meters below seafloor (mbsf), at which point basement was contacted. Roughly 1 m more penetration recovered ~5 cm of the uppermost basalt. The lower portion of the sedimentary section was re-sampled in the rotary- cored Hole 1256C from a depth of 220.10 to 245 mbsf. Sediments in Hole 1256B provide a complete sampling of the sedimentary sequence overlying the oceanic basement. The sediments are subdivided into two principal lithologic units. Unit I (0-40.6 mbsf) is clay-rich with a few carbonate-rich intervals and becomes increasingly calcareous with depth, whereas Unit II (40.6-

17 250.7 mbsf) is predominantly biogenic carbonate with some minor more siliceous and diatom- rich intervals. Bioturbation is common throughout the whole sedimentary sequence. Calcareous nannofossils are generally abundant and moderately to well preserved in all samples from Site 1256. Preservation and individual nannofossil species abundance are recorded in Table 1-1. All species considered in this report are listed in alphabetical order in the distribution chart (Table 1-1) and Appendix A. Descriptions of species and genera noted in this study and their bibliographic references are given in Perch-Nielsen (1985), Bown (1999), and Wise et al. (2004). This chapter examines the nannofossil biostratigraphy of Hole 1256B and some associated postdepositional events from the middle Miocene to Quaternary. The biostratigraphy for the sedimentary sequence immediately above the crust provides a useful tool for dating the underlying oceanic crust, because direct dating of usually-altered oceanic basalts has seldom been possible. Correlation of nannofossil datums in eastern equatorial Pacific (EEP) drill holes would indicate intra- and inter-plate variations in geologic and tectonic history.

Materials and Methods

A total of 291 samples were taken aboard JOIDES Resolution at 2 per section (at a sample spacing of 75 cm) for all cores from Hole 1256B, and smear slides were made for all samples. Preparation of smear slides followed standard techniques using Norland-61 optical adhesive as a mounting medium. Slides were examined using a Zeiss Axioskop 2 microscope under cross- polarized light, transmitted light, and phase-contrast light at 800x and 1250x magnification. A JEOL JSM 840 scanning electron microscope (SEM) was employed to take digital images for more precise species identification. Relative abundance of individual nannofossil species and overall preservation and abundance of the nannofossil assemblages were recorded in semiquantitative estimates under a magnification of 1000x for each sample. Letter codes for these estimates follow those of Arney and Wise (2003) as below. Estimates of overall nannofossil abundance were given the following letter codes: V = very abundant (>10 nannofossils/field of view). A = abundant (1-10 nannofossils/field of view). C = common (1 nannofossil/2-10 fields of view).

18 F = few (1 nannofossil/>10-100 fields of view). R = rare (1 nannofossil/>100 fields of view). B = barren (no nannofossils/>500 fields of view). The average state of preservation of the nannofossil assemblage in each sample is designated as follows: VG = very good (no evidence of dissolution and/or overgrowth; no alteration of primary morphological characteristics and specimens appear diaphanous; specimens are identifiable to the species level). G = good (little or no evidence of dissolution and/or overgrowth; primary morphological characteristics only slightly altered; specimens are identifiable to the species level). M = moderate (specimens exhibit some etching and/or overgrowth; primary morphological characteristics sometimes altered; however, most specimens are identifiable to the species level). P = poor (specimens are severely etched or exhibit overgrowth; primary morphological characteristics largely destroyed; fragmentation has occurred; specimens cannot be identified at the species and/or generic level). Relative individual species abundance estimations follow the procedure of Hay (1970) as below: V = very abundant (>10-100 specimens per field of view). A = abundant (1-10 specimens per field of view). C = common (1 specimen per 2-10 fields of view). F = few (1 specimen per 11-100 fields of view). R = rare (1 specimen per 101-1000 fields of view).

Biostratigraphy Results

The nannofossil zonation schemes proposed by Martini and Müller (1986) and Okada and Bukry (1980) for low latitudes were used as the basic zonal reference in this study. Absolute ages were assigned to all datum levels to facilitate derivation of sedimentation rates and easy comparison with other studies. Most ages were compiled from Berggren et al. (1995a, 1995b) and a few from other commonly cited sources (Table 1-2). The positions of biostratigraphic datums and their numerical ages are presented in Table 1-2.

19 Quaternary (0-1.77 Ma) The calcareous nannofossils recovered from the Quaternary are generally abundant to common and moderately preserved in Hole 1256B, and are affected to different degrees mainly by etching and fragmentation. No significant reworking of nannofossils was apparent, and the sparse occurrence of reworked discoasterids and sphenoliths in a few samples did not hinder the recognition of zonal datums. The assemblages have low diversity with ~17-21 species and are characterized by abundant to common Gephyrocapsa, Reticulofenestrata, Calcidiscus, and zonal markers for the specific zones. Zone NN21 (CN15; 0-0.26 Ma) Sediments downhole to Sample 206-1256B-1H-2, 40-42 cm (1.90 mbsf) were placed in this zone. The base of this zone is defined by the first occurrence (FO) of Emiliania huxleyi, for which recognition required confirmation under SEM due to its tiny size (2-4 μm). Zone NN20 (CN14b; 0.26-0.46 Ma) Samples 206-1256B-1H-2, 115-117 cm (2.65 mbsf), and 1H-4, 40-42 cm (4.90 mbsf), were assigned to Zone NN20, which is a “gap” zone defined by the presence of Gephyrocapsa and . absence of E. huxleyi and Pseudoemiliania ٛ artiniٛ Zone NN19 (CN14a-CN13; 0.46-1.95 Ma) -This interval lies between the last occurrences (LOs) of P. ٛ artiniٛ in Sample 206-1256B 1H-4, 115-117 cm (5.65 mbsf), and of Discoaster brouweri in Sample 3H-3, 115-117 cm (19.75 mbsf). Another three datums, the LOs of Helicosphaera sellii and Calcidiscus macintyrei, and FO of Gephyrocapsa caribbeanica, which are present in Samples 206-1256B-2H-CC (15.91 mbsf), 3H-1, 115-117 cm (16.75 mbsf), and 3H-1, 40-42 cm (16.00 mbsf), respectively, can be used to refine this interval according to Gartner (1977), Okada and Bukry (1980), and Raffi et al. (1993). Their ages are listed in Table 1-2 but not described in detail here.

Pleistocene/Pliocene Boundary The transition from the Pliocene to Pleistocene at Site 1256 is marked by an interval barren of nannofossils that spans from Samples 206-1256B-3H-2, 40-42 cm (17.50 mbsf) to 3H-3, 40- 42 cm (19.00 mbsf), and falls within Zones NN19 and NN18. The Pleistocene/Pliocene boundary is arbitrarily placed at the top of this interval (17.50 mbsf), which closely approximates the lithologic Subunit IA/IB boundary at a sediment color change at Section 206-1256B-3H-2, 38

20 cm (17.48 mbsf). This boundary, with an age of 1.78 Ma (Berggren et al., 1995b), is above the apparent LO of Discoaster brouweri (1.95 Ma) in Sample 206-1256B-3H-3, 40-42 cm (19.75 mbsf).

Pliocene (1.77-5.32 Ma) Nannofossils recorded from the Pliocene generally have a higher abundance relative to the Quaternary, except that there are several intervals barren of nannofossils, and the abundance increases with depth. The preservation of nannofossil assemblages ameliorates with increasing depth, and is closely associated with the presence and abundance of diatoms, which release silica to pore waters and thus inhibit dissolution and/or precipitation of calcite during diagenesis (Wise, 1977). The assemblages above Zone NN15 remain low in diversity and contain abundant to common reticulofenestrids, discoasterids, coccoliths, and Calcidiscus. Downhole from Zone NN15, the diversity and abundance of nannofossils increases abruptly, which is caused by the dominance and diversification of reticulofenestrids and sphenoliths. Zone NN18 (CN12d; 1.95-2.44 Ma) Samples 206-1256B-3H-3, 115-117 cm (19.75 mbsf), to 3H-5, 115-117 cm (22.75 mbsf), are assigned to Zone NN18. The interval is normally defined as the LO of D. pentaradiatus to the LO of Discoaster brouweri. Assuming a normal stratigraphic order by stratigraphic sequence, the nannofossil-barren intervals at the top and bottom of this zone can be constrained by the overlying and underlying biozones and are assigned to the combined NN19-NN18 and NN18- NN17 Zones, respectively; hence, the upper and lower boundaries of Zone NN18 are uncertain. The apparent LO of Ceratolithus rugosus is also recorded in the same sample as D. brouweri in Sample 206-1256B-3H-3, 115-117 cm (19.75 mbsf). Zone NN17 (CN12c; 2.44-2.61 Ma) As mentioned above, the upper boundary of this zone lies in a nannofossil-barren interval and thus remains uncertain. The base of this zone is recorded above Sample 206-1256B-4H-2, 40-42 cm (27.00 mbsf), based on the LO of Discoaster surculus in this sample. Zone NN16 (CN12a-b; 2.61-3.75 Ma) This zone is assigned to Samples 206-1256B-4H-2, 40-42 cm (27.00 mbsf), through 4H-4, 40-42 cm (30.00 mbsf), based on the LO of Reticulofenestra pseudoumbilica in the underlying Sample 4H-4, 115-117 (30.75 mbsf), a species defined by a minimum coccolith length of >7µm.

21 An alternative zonal marker proposed by Okada and Bukry (1980) for the base of this zone is the LO of Sphenolithus abies/neoabies, which occurs a little higher in stratigraphic position and was recorded in the lowermost sample of this zonal interval. The presence of a barren sample in the middle of this interval does not prohibit placement of the zonal boundaries. Zone NN15 (CN11; 3.75-4.8 Ma) This zone encompasses Samples 206-1256B-4H-4, 115-117 cm (30.75 mbsf), through 5H-1, 115-117 cm (35.75 mbsf), which directly overly the LO of Amaurolithus primus in 5H-2, 40-42 cm (36.50 mbsf), an equivalent zonal marker proposed by Okada and Bukry (1980) to replace the rare Amaurolithus tricorniculatus. The latter only occurs in Sample 206-1256B-5H-2, 115- 117 cm (37.25 mbsf). Zone NN14-NN13 (CN10c; 4.8-5.04 Ma) Sediments between Samples 206-1256B-5H-2, 40-42 cm (36.50 mbsf) and 5H-3, 115-117 cm (38.75 mbsf), an interval directly above the LO of Ceratolithus rugosus in 5H-4, 40-42 cm (39.50 mbsf), are assigned to this combined zone. Another useful datum, the LO of Ceratolithus acutus, occurs in Sample 206-1256B-5H-2, 115-117 cm (37.25 mbsf). No attempt was made to determine the FO of Discoaster asymmetricus, a datum defining the boundary between Zones NN14 and NN13, due to the intermittent occurrence of this species. Zone NN12 (CN10a-b; 5.04-5.34 Ma) This zone is recorded in the interval from Samples 206-1256B-5H-4, 40-42 cm (39.50 mbsf), to 5H-5, 115-117 cm (41.75 mbsf), where both C. rugosus and Discoaster quinqueramus are absent. The FO of D. acutus falls within this zone, which helps in the recognition of this “gap” zone.

Pliocene / Miocene Boundary The Pliocene/Miocene boundary in terms of calcareous nannofossils falls within Subzone CN10a or Zone NN12 (Perch-Nielsen, 1985). The FO of D. acutus closely approximates this epoch boundary, which lies just above the extinction of D. quinqueramus, a well-documented datum level in most sections.

22 Miocene (5.32 -23.80 Ma) Nearly all the Miocene samples examined contain very abundant and diversified nannofossil assemblages. These assemblages are generally moderately to well preserved but deteriorate downhole, due mainly to overgrowth as well as dissolution. Similar to preservation in the overlying Pliocene assemblages, the preservation depends to a great extent on the presence and abundance of diatoms. The assemblages consist of highly diversified and predominant reticulofenestrids, discoasterids, sphenoliths, coccoliths, and Calcidiscus. Zone NN11 (CN9; 5.6-8.6 Ma) This zone covers the stratigraphic range of D. quinqueramus, which was observed between Samples 206-1256B-5H-6, 40-42 cm (42.50 mbsf), and 9H-6, 40-42 cm (80.50 mbsf); in the latter sample the FO of Discoaster berggrenii was documented. The FO of Amaurolithus primus, located in Sample 206-1256B-7H-7, 40-42 cm (63.00 mbsf), is used to subdivide this thick zone into upper NN11b and lower NN11a Subzones. Two other events, the LO and FO of Amaurolithus amplificus in Samples 206-1256B-6H-3, 40-42 cm (47.50 mbsf), and 7H-2, 40-42 cm (55.50 mbsf), respectively, can be used with caution due to its generally rare occurrence to further refine Subzone NN11b. The LO and FO of Reticulofenestra rotaria, a circular placolith with relatively large central opening and a diameter of 5-7 µm, recorded in Samples 206-1256B-6H-6, 115-117 cm (52.75 mbsf), and 7H-CC (63.39 mbsf) between Chrons C3An.2n (t) (51.60 mbsf) and C4n.1n (o) (71.56 mbsf) (Shipboard Scientific Party, 2003c), have the stratigraphic potential to refine Subzone NN11b. Assuming a constant sedimentation rate, derived independently from biostratigraphy and magnetostratigraphy, the FO is calculated at 7.18 and 7.03 Ma, and its LO at 6.32 and 6.34 Ma, respectively. The difference in the age estimates probably arises from sedimentation rate variation and sample spacing for biostratigraphy; the latter issue in sample spacing could be resolved by further closer sampling. Considering the stratigraphic distribution of R. rotaria, A. primus, and A. amplificus in the sequence (Table 1-1), the ages for the FO and LO of R. rotaria are assigned to 7.18 and 6.32 Ma, respectively. Taking into consideration of the differences in time scales used in both studies, these assignments are reasonably consistent with those calibrated by Wei (1995). Zone NN10 (CN8; 8.6-9.4 Ma)

23 This zone is assigned to Samples 206-1256B-9H-6, 115-117 cm (81.25 mbsf), through 10H- 2, 40-42 cm (84.00 mbsf), the base of which is defined by the LO of Discoaster hamatus in Sample 10H-2, 115-117 cm (84.75 mbsf). This datum coincides with the LO of Discoaster bollii in this hole. The presence of Discoaster neorectus aids in identifying this gap zone. Zone NN9 (CN7; 9.4-10.38 Ma) This zone spans the stratigraphic range of D. hamatus, and encompasses Samples 206- 1256B-10H-2, 115-117 cm (84.75 mbsf), through 12H-2, 115-117 cm (103.75 mbsf). Shipboard observation of this marker in Sample 206-1256B-10H-4, 100 cm (87.60 mbsf), is consistent with this assignment. Because the upper boundary is populated by the rare and sparse occurrence of this species, and there exist samples bearing very few or no nannofossils, the precise placement requires much effort. It is worth mentioning that the occurrences of this zonal marker in Samples 206-1256B-9H-4, 40-42 cm (77.50 mbsf), and 115-117 cm (78.25 mbsf), are considered to be reworked, the significance of which will be discussed in the “Discussion” section. The LO of Catinaster coalitus recorded in Sample 206-1256B-11H-3, 115-117 cm (95.75 mbsf), can be used to subdivide this zone. The intermittent occurrence of Discoaster neohamatus in the lower part of its range at this site renders the application of its FO datum impossible. Zone NN8 (CN6; 10.38-10.9 Ma) This zone lies between the Fos of D. hamatus and C. coalitus, the latter datum being observed in Sample 206-1256B-12H-4, 40-42 cm (106.00 mbsf). The LO of Coccolithus miopelagicus approximates the top of this zone. Zone NN7 (CN5b; 10.9-11.8 Ma) Zone NN7 is recorded by the FO of C. coalitus downhole to the FO of Discoaster kugleri in Sample 206-1256B-17H-1, 40-42 cm (149.00 mbsf). The preservation of nannofossils deteriorates with depth from the middle of this zone, which makes the identification of discoasters difficult. D. kugleri, however, can still be distinguished from Discoaster sanmiguelensis and other discoasterids in overgrown specimens by the diagnostic six short rays and a broad, flat central area. Its LO recorded in Sample 206-1256B-14H-CC (130.05 mbsf) provides a datum to further subdivide this thick interval. Zone NN6 (CN5a; 11.8-13.6 Ma) The lowermost biostratigraphic zonal event in the section, the LO of Sphenolithus heteromorphus, the specimens of which are overgrown but recognizable by their diagnostic

24 bright apical spines and butterfly-shaped bases, marks the base of Zone NN6 in Sample 206- 1256B-25X-1, 115-117 cm (214.15 mbsf). Here Cyclicargolithus floridanus, the last continuous occurrence (LCO) of which is in Sample 206-1256B-23X-CC (202.77 mbsf), provides a further datum to subdivide Zone NN6. This additional datum provides a sound biostratigraphic correlation line in equatorial areas (Raffi et al., 1995). Zone NN5 (CN4; 13.6-15.6 Ma) Helicosphaera ampliaperta, the LO of which marks the base of Zone NN5, was absent from all samples including the lowermost Sample 206-1256B-28X-CC (245.83 mbsf). This indicates that the basaltic basement underlying the sedimentary sequence is younger than 15.6 Ma (age of the LO of H. ampliaperta). The rare and sparse, but distinctive, onset of R. pseudoumbilica is in Sample 206-1256B-26X-3, 40-42 cm (226.10 mbsf), and provides a useful datum to assess the age of basement. Assuming a constant sedimentation rate of 39.1 m/m.y., based on datums from below 100 mbsf, a basement age of ~14.5 Ma is implied (Fig. 1-2), consistent with the assignment to Zone NN5 of the middle Miocene. This age is reasonably consistent with the age of 15.1 Ma inferred from magnetic anomalies (Shipboard Scientific Party, 2003c).

Discussion

With two-thirds of the cores taken by APC as well as good sediment recovery via rotary coring, Leg 206 has added considerably to the sediment recovery for this region. The high core recovery (89%) provides an excellent opportunity to study environmental, oceanographic, and biotic changes from the middle Miocene to the Holocene in the equatorial Pacific. Sedimentation rates especially were generally high in those critical time intervals, which is ideal for high- resolution biostratigraphic and paleoceanographic studies.

Sedimentation History Linear sedimentation rates, uncorrected for compaction, were calculated based on 28 nannofossil age estimates in Table 1-2 for the sedimentary sequence of Hole 1256B and plotted with best-fit lines. Sedimentation rates vary from 6.1 to 39.1 m/m.y. (Fig. 1-2A). Calculated sedimentation rates are high in the middle Miocene (39.1 m/m.y.), decrease drastically in the lower upper Miocene (8.4 m/m.y.), and recover somewhat in the uppermost Miocene (13.2

25 m/m.y.), reach the lowest point in the Pliocene (6.1 m/m.y.), but increase thereafter (11.8 m/m.y.). In order to illustrate changes in depositional environments and processes through time at Site 1256, mass accumulation rates (MARs) have been calculated for bulk sediments and the main sedimentary components such as calcium carbonate, terrigenous material, and biogenic silica (Fig. 1-3). Calcium-carbonate MARs show an abrupt transition at ~11 Ma, which coincides with a “diatom mat” (i.e., a series of diatom-rich laminae that occur at this horizon throughout the region; see Kemp and Baldauf, 1993). Carbonate MARs are high except for a major drop at ~12.9 Ma during the middle Miocene, remain extremely low from ~10.2 to 9.1 Ma, and recover somewhat from 9 to ~ 5 Ma, then stay particularly low thereafter. This pattern of variations in sedimentation was also observed at Sites 844-853 drilled during ODP Leg 138 (Fig. 1-1). Although there is a difference in proximity to continents and in water depths among sites located in Guatemala basin (844, 845, and 1256), sedimentation rates positively covary (Fig. 1-2B), and sedimentary thickness during a specific time interval shows little variation (Fig. 1-4). This suggests that these sites have experienced a virtually same sedimentation history and terrestrial influences have remained a minor factor from middle Miocene through Quaternary. This, together with the fact that positions of the epoch boundaries are closely related to water depths (Fig. 1-4), reflects virtually no intra-plate variation in geological and tectonic settings on Cocos plate. The nature of deposition in EEP has been influenced mainly by surface-water productivity. Sites on the Cocos and Nazca plates are separated from the west coast of Central America by the Middle American Trench, which traps most terrestrial sediments shed from the continent; and sites on the Pacific plate are far removed from continents. The sedimentary sources, therefore, are mainly biogenic calcite and silica, which determine the sedimentation rates, e.g., at Site 1256

(Fig. 1-3). This is further confirmed by the close correlation between CaCO3 MARs and Ba/Ti ratios, chemical proxies for biogenic productivity (Shipboard Scientific Party, 2003c). The distribution pattern of surface-water productivity today helps explain the spatial and temporal variations in EEP. The V-shape profile in the longitudinal transect along Sites 848 through 853 clearly reflects diminishing productivity away from the equator (Fig. 1-4), as seen today (Paytan et al., 1996). Sites with substantially higher sedimentation rates are located under

26 highly productive waters, e.g., Sites 849-851 under the equatorial divergence, and Sites 846-847 close to the Galápagos Islands under strong influence of the Peru upwelling system. These sites have been remaining under this situation since middle Miocene due to a moving path nearly parallel to the equator (Pisias et al., 1995). The high biological productivity fuels the high sedimentation rates. Conversely, the diminishing productivity has resulted in a decrease in sedimentation rates as sites on the Cocos plate (844, 845, and 1256) move away from the equator. The high sedimentation rates in the middle Miocene can be attributed to high productivity and good preservation while these sites were near the paleoequator on young, shallow seafloor. This is indicated by a nannofossil assemblage showing little dissolution at Site 1256. The lowest sedimentation rates (Pliocene), however, likely have resulted from intense dissolution superimposed on reduced paleoproductivity as indicated by the poor to moderate preservation of nannofossils and nannofossil-barren intervals at Site 1256. This probably resulted from a combination of sea-floor subsidence and a shoaling of the calcite compensation depth (CCD) associated with a change in chemistry of deep waters and/or surface carbonate production.

Middle/Late Miocene Carbonate Crash The major middle-late Miocene carbonate shift recorded at Site 1256 is referred to as “the carbonate crash” by Lyle et al. (1995) and Farrell et al. (1995). This event has been widely recorded in the Pacific and Atlantic oceans and Caribbean Sea (Berger, 1970; King et al., 1997; Roth et al., 2000) and provides a seismic reflector for long-range correlation (Mayer et al., 1986; Bloomer et al., 1995). Lyle et al. (1995) and Farrell et al. (1995) attributed the carbonate crash to enhanced dissolution associated with changes in bottom-water characteristics, rather than with changes in productivity. Emerson and Bender (1981) and Archer (1991a, 1991b), however, have demonstrated that under proper conditions, increases in productivity can lead to enhanced dissolution through acids produced by the degradation of organic matter. Lyle et al. (1995) cited a decrease rather than increase of organic carbon to support their argument, which seems plausible because such a decrease in organic carbon can also be caused by enhanced degradation resulting in higher acid production. In such a case the organic carbon cannot be used as proxy of productivity. Several mechanisms have been proposed for the carbonate crash. Even though there is an apparent correlation of the carbonate crash and a eustatic sea-level drop (Haq et al., 1987),

27 Peterson et al. (1992) showed that globally, the middle-late Miocene sea-level drop is associated with deepening of the CCD and enhanced carbonate preservation. This is supported by Berger’s (1970) basin-shelf fractionation model. Other causative mechanisms proposed, such as the constriction of the Panama Gateway (Lyle et al., 1995) and the initiation of North Atlantic Deep Water (NADW) (Woodruff and Savin, 1989), account for the crash by enhanced corrosiveness of bottom water imported into this region. The arguments presented above clearly show that the middle/late Miocene carbonate crash is still not resolved in terms of the interplay of dissolution and production. Dissolution seems to play a certain role associated with large variations in the CCD, with nearly complete carbonate dissolution during brief intervals when the CCD shoaled by as much as 1400 m (van Andel et al., 1975a, b). However, the cause of enhanced dissolution is complex, but probably relates to (1) tectonic movement associated with the Panama Gateway and a slowdown of sea-floor spreading, as indicated by the magnetic anomaly pattern in Figure 1; and (2) phytoplankton productivity associated with ocean circulation patterns. Evidence from calcareous nannofossils and diatoms tends to attribute it to surface-water process. Prior to the carbonate crash, all major reductions in carbonate resulted from deposition of laminated diatom oozes (Kemp, 1995), which is believed to be controlled by surface process, instead of dissolution at depth (Kemp et al., 1995). At Site 1256, the calcareous nannofossil assemblages exhibit their best preservation just prior to the carbonate crash nadir. The same is true at Site 846 (Raffi and Flores, 1995). This proposal is consistent with the study in terms of stable isotopes (this volume).

Downslope Transport of Sediments Calcareous nannofossils are susceptible to reworking because of their tiny size, hence reworked fossils often pose problems in biostratigraphy. Some of them with a short range but that are clearly out of place (e.g., D. hamatus), together with lithologic information, however, provide an opportunity to study episodes of redeposition. For example, downhole investigation of nannofossils found D. quinqueramus and D. hamatus together in Samples 206-1256B-9H-4, 40-42 cm (77.50 mbsf), and 115-117 cm (78.25 mbsf), which caused confusion in interpreting the biostratigraphy. Upon close examination of more samples, the isolated occurrences of D. hamatus in these two samples seem more likely to be reworked, and hence its actual LO should

28 be much lower, which is supported by the sudden sedimentary color change. Based on the criteria of reworked fossils and abrupt lithologic change, three episodes of redeposition were recognized, and the timing for these were obtained from the age-depth plot assuming constant sedimentation rates for the specific interval. Episode 1: This reworking episode is recognized at Section 206-1256B-5H-1, 108 cm (35.68 mbsf), which dates to ~4.7 Ma. It is recognized by the common reoccurrence of D. berggrenii and D. quinqueramus from Zone NN11 along with Sphenolithus moriformis in Samples 206-1256B-4H-CC (35.23 mbsf), and 5H-1, 40-42 cm (35.00 mbsf). Note the deeper depth of Sample 4H-CC relative to 5H-1, 40-42 cm, because of the >100% core recovery for Core 4H. A sharp uphole change in sediment color from dark brown to gray brown and reduced bioturbation occur across the boundary. Episode 2: Against the normal background deposition in Zone NN11, characterized by the presence of D. quinqueramus and D. berggrenii, reworked specimens of D. hamatus and D. bollii stand out in Samples 206-1256B-9H-4, 40-42 cm (77.50 mbsf), and 115-117 cm (78.25 mbsf). The conspicuous sudden reduction in bioturbation upwards and sedimentary color change from green gray to dark green co-occurs right between Sections 9H-4 and 9H-5, and the boundary between these is placed at Section 206-1256B-9H-4, 150 cm (78.60 mbsf), with an age of ~ 8.3 Ma. Episode 3: The oldest redepositional event is recognized at Section 206-1256B-11H-6, 95 cm (100.05 mbsf), and is dated at 10.7 Ma. It is signaled by reworked specimens of C. abisectus (Oligocene to early Miocene age) and an abrupt lithologic change. Upward across this boundary, there is not only a sudden change in sediment color from dark green gray to light green gray, but also a loss of intense bioturbation. These redepositional events provide cogent evidence for downslope processes active on the EPR throughout Neogene. It is worth noting that Episode 2 has been recorded in Knüttel’s 1986 study on the EPR as well as in many ocean basins (Ciesielski and Wise, 1977; Adams et al., 1977). Knüttel (1986) suggested that a Messinian sea-level drop triggered intensified bottom current activity, whereas Rea and Janecek (1986) alternatively interpreted this as tectonic instability associated with possible increased seismic activity. By either mechanism, however, turbidity currents have produced a rapid accumulation of sediments as indicated by the sharp sedimentary contacts and sharply reduced bioturbation.

29 Conclusions

The primary objective of this paper has been to refine middle Miocene to Quaternary nannofossil biostratigraphy of Site 1256 in the Guatemala Basin on Cocos plate crust to compare the data with those from other sites in this basin. 1. The middle Miocene through Quaternary sedimentary sequence at Site 1256 appears to be complete in terms of the calcareous nannofossil assemblage succession. A total of 16 zones/combined zones are recognized, and nearly all zonal boundaries are located within a quarter section of a core. 2. Age estimates for two nannofossil datums, 7.18 Ma for the FO, and 6.32 Ma for the LO of R. rotaria, respectively, were derived from biostratigraphy and magnetostratigraphy. 3. Linear sedimentation rates were calculated using 28 nannofossil age estimates. Rates are high in the middle Miocene, decrease drastically in the lower upper Miocene and recover somewhat in the uppermost Miocene, reach the lowest point in the Pliocene, but increase thereafter. This pattern of variations is similar to that observed at nearby ODP sites in the Guatemala Basin. The middle-upper Miocene carbonate crash was observed. The crash is believed to have been caused by enhanced dissolution that resulted from lowered productivity as well as the initiation of NADW and the constriction of the Panama Gateway. These similar sedimentary patterns, together with the close correlation of epoch boundaries at Site 1256 to those published for other nearby DSDP/ODP sites, suggest virtually no intra-plate variations on the Cocos plate in terms of depositional history. 4. Reworked nannofossils, together with sedimentation changes indicated by significant color change, were used to investigate redeposition events, and three episodes were recognized. The Messinian sea-level drop and increased seismic activity may have contributed to intensified bottom current activity. Turbidity currents likely produced these episodes of redeposition as indicated by a sharp change in sediment color and a sharp reduction in bioturbation.

Acknowledgements

We wish to thank all the participants of ODP Leg 206 for their contributions relevant to this paper, and the members from the Calcareous Nannofossil Laboratory at Florida State University

30 for their helpful comments and assistance. We thank Drs. Gary Acton, José-Abel Flores, and Koji Kameo for helpful and constructive reviews. This research used samples and/or data provided by the Ocean Drilling Program (ODP). ODP is sponsored by the U.S. National Science Foundation (NSF) and participating countries under management of Joint Oceanographic Institutions (JOI), Inc. Funding for this research was provided by a United States Science Advisory Committee (USSAC) grant to both authors.

Figure 1-1. Backtracked path for Site 1256, Leg 206, and locations of Leg 138 sites. The backtracking is shown at 1-Ma increments, with the last point in the path indicating the location of the site when the basement formed (after Shipboard Scientific Party, 2003c, and Pisias et al., 1995).

Figure 1-2A. Linear sedimentation rates as constrained by the biostratigraphy and magnetostratigraphy, and a best-fit linear sedimentation rate model. The linear rates are shown extrapolated to basement. Nannofossil datum age estimates and depths are presented in Table 1-2. FO = first occurrence, LO = last occurrence. Magnetostratigraphy data are from Shipboard Scientific Party (2003).

Figure 1-2B. Calcareous nannofossil datum age-depth plots for eastern equatorial Pacific drill holes. Data for Sites 844-853 are from Raffi and Flores (1995).

Figure 1-3. Relationship of mass accumulation rates between bulk sediments and the main sedimentary components. Note the change of scale for CaCO3 in the 0-9 Ma interval.

Figure 1-4. Correlation of nannofossil datums in EEP drill holes that reflects variations in sedimentation rates over time. Data for Sites 844-853 are from Raffi and Flores (1995). Heavy lines approximate epoch boundaries. The numerical age for each correlation line is listed in chronological order in the lower right box.

36 Table 1-1. Calcareous nannofossil distribution chart in ODP Hole 1256B. Notes: Preservation: VG = very good, G = good, M = moderate, P = poor. Abundance: V = very abundant, A = abundant, C = common, F = few, R = rare, B = barren; abundance shown in lower case indicates reworking. See “Materials and Methods” for explanation.

38 Table 1-1 – continued

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59 Table 1-1 – continued

60 Table 1-2. Age estimate, depth, and zonation of calcareous nannofossil datum levels used in this study. Notes: FO = first occurrence, LO = last occurrence, LCO = last continuous occurrence, Ref. = reference. The zonation codes are after Martini (1971) and Martini and Müller (1986) (NN zones), and Okada and Bukry (1980) (CN zones). 1: Berggren et al. (1995a); 2: Raffi and Flore (1995); 3: Shackleton et al. (1995); 4: Berggren et al. (1995b); 5: Backman and Raffi (1997).

Zone Core, section, interval (cm) Depth (mbsf) Age Biohorizons Ref. (base) Top Bottom Top Bottom Mean (Ma) 206-1256B- 206-1256B- FO Emiliania huxleyi NN21/CN15 1H-2, 40-42 1H-3, 40-42 1.90 3.40 2.65 0.26 1 LO Pseudoemiliania lacunosa NN20/CN14b 1H-4, 40-42 1H-4, 115-117 4.90 5.65 5.28 0.46 1 LO Helicosphaera sellii 2H-7, 40-42 2H-CC 15.47 15.91 15.69 1.47 1 LO Calcidiscus macintyrei NN19b 3H-1, 40-42 3H-1, 115-117 16.00 16.75 16.38 1.59 1 LO Discoaster brouweri NN19a/CN13 3H-1, 115-117 3H-3, 115-117 16.75 19.75 18.25 1.95 1 LO Discoaster pentaradiatus NN18/CN12d 3H-5, 115-117 3H-7, 40-42 22.75 25.00 23.88 2.44 2 LO Discoaster surculus NN17/CN12c 4H-1, 115-117 4H-2, 40-42 26.25 27.00 26.63 2.61 2 LO Sphenolithus abies/neoabies 4H-3, 40-42 4H-4, 40-42 28.50 30.00 29.25 3.60 1 LO Reticulofenestra pseudoumbilicus NN16/CN12a 4H-4, 40-42 4H-4, 115-117 30.00 30.75 30.38 3.75 1 LO Amaurolithus spp. (A. primus) CN11a 5H-1, 115-117 5H-2, 40-42 35.75 36.50 36.13 4.80 1 LO Ceratolithus acutus 5H-2, 40-42 5H-2, 115-117 36.50 37.25 36.88 4.99 3 FO Ceratolithus rugosus NN13/CN10c 5H-3, 115-117 5H-4, 40-42 38.75 39.50 39.13 5.04 2 FO Ceratolithus acutus CN10b 5H-4, 115-117 5H-5, 40-42 40.25 41.00 40.63 5.34 1 LO Discoaster quinqueramus NN12/CN10a 5H-5, 115-117 5H-6, 40-42 41.75 42.50 42.13 5.60 4 LO Amaurolithus amplificus 6H-2, 115-117 6H-3, 40-42 46.75 47.50 47.13 5.90 4 FO Amaurolithus amplificus 7H-3, 40-42 7H-3, 115-117 57.00 57.75 57.38 6.60 4 FO Amaurolithus primus NN11b/CN9b 7H-7, 40-42 7H-CC 63.00 63.39 63.20 7.20 4 FO Discoaster berggrenii NN11a/CN9a 9H-6, 40-42 9H-6, 115-117 80.50 81.25 80.88 8.60 4 LO Discoaster hamatus NN10/CN8a 10H-2, 40-42 10H-2, 115-117 84.00 84.75 84.38 9.40 4 LO Catinaster coalitus 11H-3, 40-42 11H-3, 115-117 95.00 95.75 95.38 9.69 5 FO Discoaster hamatus NN9/CN7a 12H-2, 115-117 12H-2, 115-117 103.75 103.75 103.75 10.38 3 LO Coccolithus miopelagicus 12H-3, 40-42 12H-3, 40-42 104.50 104.50 104.50 10.80 4 FO Catinaster coalitus NN8/CN6 12H-4, 40-42 12H-4, 115-117 106.00 106.75 106.38 10.90 4 LO Discoaster kugleri 14H-7, 40-42 14-CC 129.50 130.05 129.78 11.50 4 FO Discoaster kugleri NN7/CN5b 17H-1, 40-42 17H-1, 115-117 149.00 149.75 149.38 11.80 4 LCO Cyclicargolithus floridanus 23X-6, 115-11723X-CC 202.35 202.77 202.5613.19 2 LO Sphenolithus heteromorphus NN6/CN5a 25X-1, 40-42 25X-1, 115-117 213.40 214.15 213.78 13.60 4 FO Reticulofenestra pseudoumbilicus 26X-3, 40-42 26X-CC 226.10 226.71 226.41 13.95 3

CHAPTER 2

TAXONOMIC NOTE: DISCOASTER STELLULUS GARTNER, 1967, EMENDED

Introduction

Discoaster stellulus was first described by Gartner (1967) from the Middle Miocene Lengua Formation, Trinidad. The original description, however, was largely based on the distal view of the asterolith, while the proximal view was not depicted in electron micrographs and, therefore, essentially remained undescribed. This may have caused misplacement of specimens in proximal view in a different species, hindering true identification and, consequently, application of this species in biostratigraphy.

A calcareous nannofossil biostratigraphic study on the diatomaceous materials from Hole 1256B, Ocean Drilling Program (ODP) Leg 206 to the eastern equatorial Pacific, yielded abundant well-preserved specimens of D. stellulus. Both views of this asterolith were extensively examined using light and electron microscopy. A mobile-mount technique was applied to observe multiple views of same specimens. All of these allowed a redescription and reillustration of this species, the definition of which is emended herein.

Materials and Methods

The studied materials are dark greenish-gray nannofossil-diatom oozes from a Middle Miocene to Quaternary carbonate-rich section recovered during ODP Leg 206 to the Guatemala

62 Basin. Sixteen zones/combined zones and 28 datums were recognised in the sedimentary sequence in Hole 1256B, according to the biostratigraphic schemes of Martini (1971), Martini & Müller (1986) and Okada & Bukry (1980). The nannofossil assemblages are characteristic of low latitudes, with abundant Gephyrocapsa, Discoaster and Sphenolithus, and are, in general, moderately- to well-preserved, depending on nannofossil abundance and the presence of diatoms. The two consecutive samples containing Discoaster stellulus were assigned to Zone NN11d, based on the co-occurrence of Discoaster quinqueramus and Amaurolithus amplificus (Jiang & Wise, in press). Preparation of smear-slides followed the standard technique (Bown, 1998), using Norland 61 optical adhesive as a permanent mounting medium. Selected samples were also prepared and examined in a mobile mount. Unprocessed sediments were smeared on a glass coverslip, dried, and then mixed with Norland 61 optical adhesive as evenly as possible, and the coverslip placed on a glass slide. The slide was not cured under ultraviolet light so that specimens could be rotated by gently moving the coverslip to create flows within the viscous mounting medium. This allowed observation of both sides of a specimen, as well as a side-view of the same specimen. Slides were examined using a Zeiss Axioskop II microscope under cross-polarised light (XPL), transmitted light (TL), and phase-contrast light (PC) at 1000x magnification. A JEOL JSM 840 scanning electron microscope (SEM) was employed to observe fine-scale structures. All materials are deposited in the collections of the Calcareous Nannofossil Laboratory at the Department of Geological Sciences, Florida State University (FSU), Tallahassee, Florida, USA.

Systematic Palaeontology

Family DISCOASTERACEAE Tan, 1927 Genus Discoaster Tan, 1927 Discoaster stellulus Gartner, 1967, emend. Jiang & Wise, 2006 Pl.2.1, figs 1-22

63 Emended diagnosis: Six-rayed asterolith with elongated, diamond-shaped depressions on the proximal side and prominent parallel-sided ridges along the rays on the distal side. Original description: Asterolith, usually with 6 short rays; prominent parallel-sided ridges radiate from center to near tip of each ray. Emended description: The asterolith consists of 6 rays, each with a prominent parallel-sided ridge on the distal side, whereas the arms display elongate, diamond-shaped depressions on the proximal side. The arms are stubby, tapering abruptly to a single point or a notched bifurcation. The bifurcations have short, pointed limbs that form an obtuse angle. The central area is very large relative to the entire size of the asterolith, and features a large Grayed knob. The knob on the distal side is more robust and larger than on the proximal side. The asterolith appears dim in transmitted light except that the ridges along the arms are brighter and stand out. In overgrown specimens, all of these features axe more distinctive under both transmitted and polarised light. Differentiation: D. stellulus resembles in outline D. adamanteus Bramlette & Wilcoxon, 1967, which also has short, stubby, gradually tapering arms. However, in addition to bearing a 6-rayed knob on both sides, D. stellulus has diamond-shaped depressions on the proximal side and prominent parallel-sided ridges along the rays on the distal side, which differentiate this species from other Neogene discoasters. Dimensions: 4-8μm (holotype = 6.4μm). Occurrence: D. stellulus was common to abundant between the last uphole occurrences of Amaurolithus amplificus and Amaurolithus primus within Zone NN11d (uppermost Miocene) in our study section from Hole 125613. Specifically, it was observed in Samples 2061256B-6H-2, 115-117cm (46.75mbsf) and -6H-2, 4042cm (46.Oombsf). However, Gartner (1967) originally described the taxon from the Globorotalia menardii Planktonic Foraminiferal Zone, of the Lengua Formation of Trinidad (Sample C887M), which corresponds to the lowermost Upper Miocene Discoaster hamatus Zone of Bramlette & Wilcoxon (1967), or to Zone NN9 in the Martini (1971) zonation. Therefore, our observations extend the range of the species from Zone NN9 up into Subzone NN11b, although no occurrences of this species were observed in zones/subzones other than NN11d in this present study.

64 Acknowledgements

We wish to thank Kim Riddle of the FSU Biology Unit 1 SEM laboratory for helpful assistance. The Calcareous Nannofossil Laboratory at FSU Department of Geological Sciences provided lab facilities. We are grateful to Drs Isabella Raffi, Jeremy Young and Jackie Lees for constructive and helpful comments that greatly improved this manuscript. The samples were provided by the Ocean Drilling Program (ODP). ODP is sponsored by the US National Science Foundation (NSF) and participating countries under management of Joint Oceanographic Institutions (JOI), Inc. Funding for this research was provided by a United States Science Advisory Committee (USSAC) grant to SJ and SWW. General lab support was provided by NSF-OPP 0126218 to SWW.

SEM, proximal view SEM, distal view SEM, distal view SEM, proximal view

SEM, distal view SEM, oblique distal view; same SEM, proximal view SEM, oblique proximal view; same specimen as fig.9 specimen as fig.11

PC, distal view TL, proximal view; same PC, distal view PC, distal view; same PC, proximal view, same specimen as fig.13 specimen as fig.15 specimen as fig.15

TL, distal view TL, proximal view TL, side view TL, side view, mobile mount PC, oblique view Plate 2-1. Discoaster stellulus Gartner, 1967, emend. Jiang & Wise, 2006. Sample ODP 206- 1256B-6H-2, 115-117cm. Scale-bars = 1μm (1-12) or 2μm (13-22). SEM = scaning electronic microscope; TL= transmitted light; PC = phase contrast.

CHAPTER 3

CAUSE OF THE MIDDLE/LATE MIOCENE CARBONATE CRASH: DISSOLUTION OR LOW PRODUCTIVITY?

Abstract

The late/middle Miocene “carbonate crash”, a sharp decrease in carbonate mass- accumulation rates in the eastern and central equatorial Pacific as well as the Caribbean region, has previously been considered only a dissolution event associated with changes in global ocean chemistry, which is in turn believed to be tied to the production of the North Atlantic Bottom Water and/or ventilation via the Panama Seaway. δ13C data in the bulk-isotope record from Ocean Drilling Program Site 1256 show a close

parallel with CaCO3 mass accumulation rates (MARs) in the 5-14 Ma interval (correlation coefficient = 0.87), suggesting a relationship likely coupled to surface-water productivity by calcite-secreting organisms. The decoupling between δ13C and MARs after 5 Ma probably indicates a dominance of dissolution over carbonate production. Therefore, the coincidence in 13 δ C excursions with the stages of sharp reduction in CaCO3 MARs during the carbonate crash, points to a causative mechanism induced by surface circulation-induced low productivity. We speculate that the major middle/late Miocene sea-level drop may have caused the complete closure of the Indonesian Seaway. The blockage of Indonesian Throughflow would have resulted in a piling-up of surface warm water in the west Pacific, thereby strengthening the Equatorial Undercurrent system. The eastward spread of this nutrient-poor water then warmed sea-surface temperature and reduced upwelling in the central and eastern Pacific, reducing in turn biological productivity of phytoplankton. A coincident reduction in Central America and

67 circum-Caribbean volcanism plus the deflection of the delivery of volcanic ash as a result of the then prevailing southeastern trade winds across the equator further deprived these regions of trace-element nutrients, which added to lowered surface-water carbonate production. Surface- water warming and reduced upwelling is documented by negative excursions in δ18O values. The reduction in carbonate supply to the deep waters caused a rapid shoaling of the carbonate

compensation depth and triggered the carbonate crash. The close correlation between CaCO3 mass-accumulation rates and biological productivity suggests that the carbonate crash is best characterized as a low productivity event. Key Words: carbonate crash, bulk stable isotopes, paleoproductivity, paleoceanography

Introduction

“Carbonate crash” is the term applied by Lyle et al. (1995) and Farrell et al. (1995) to a major middle/late Miocene carbonate shift in tropical regions characterized by a dramatic reduction of calcium-carbonate content in sediments and poor preservation of calcareous microfossils. This phenomenon has been widely documented in the equatorial Pacific, Atlantic, and Indian Oceans (Lyle, 2003; King et al., 1997; Peterson et al., 1992; respectively) as well as in Caribbean Sea (Roth et al., 2000), and provides a seismic reflector for long-range correlation (Mayer et al., 1986; Bloomer et al., 1995). Several hypotheses have been advanced to account for the carbonate crash. Mechanisms recently proposed attributed it to enhanced dissolution (rather than reduced surface-water productivity of calcite-secreting organisms) associated with changes in deep-water circulation and shoaling of carbonate compensation depth (CCD) and/or lysocline (e.g., Lyle et al., 1995; Farrell et al., 1995; Roth et al., 2000). Alternatively, the crash may have resulted from a biologic bloom (Theyer et al., 1985), i.e., decomposition of organic matter from increases in surface productivity can result in enhanced dissolution of carbonate through acid production where the ratio of organic carbon to carbonate is high in falling debris (Emerson and Bender, 1981; Archer, 1991a, 1991b). Dilution by terrigenous or other non-carbonate sediments could also have caused carbonate reduction in the sediments (Keller and Barron, 1983; Diester-Haass et al., 2004). This scenario involves changes in sedimentation patterns probably driven by climate change and/or tectonic changes in basin configurations or closure of oceanographic pathways. A contemporary

68 global sea-level drop (Haq et al., 1987) was excluded because it is believed that the middle/late Miocene sea-level drop is associated with a deepening of the CCD and enhanced carbonate preservation (Peterson et al., 1992; Berger, 1970). The carbonate crash recorded in the Caribbean and major world oceans, however, seems to point to a common cause as indicated by the comparable nature and time-overlap of its occurrences, a cause related to processes affecting the global carbonate and carbon budget. A determination of the role that carbonate dissolution, production, and dilution have played would help unravel the cause of the crash. We present here an oxygen and carbon stable isotopic record and its correlation with carbonate mass accumulation rates (MARs) in sediments from Ocean Drilling Program (ODP) Site 1256 located in the eastern equatorial Pacific (Fig. 3-1). We then discuss the possible cause of the middle/late Miocene carbonate crash in light of these data.

Background

Study Area Site 1256 (6°44.2’N, 91°56.1’W; Fig. 3-1), drilled during ODP Leg 206, lies in 3635 m of water depth in the Guatemala Basin on Cocos-plate crust that formed at ~15 Ma on the eastern flank of the East Pacific Rise (EPR) when the site experienced a super-fast spreading rate (Wilson, 1996). This site formed at an equatorial latitude within the equatorial high-productivity zone and initially experienced very high sedimentation rates (39.1 m/m.y.; e.g., Jiang and Wise, this volume). The high core recovery (89% with two-thirds of the cores taken by the advanced piston corer) and high sedimentation rates through critical time intervals provide an excellent opportunity for high-resolution paleoceanographic study. The modern oceanographic setting in the equatorial world oceans is illustrated in Figure 1. This circulation pattern is a product of tropical atmospheric circulation and the Coriolis Effect across the equator. Because the southeastern trade winds are stronger than the northeastern antithesis, they converge north of the equator to form the Intertropical Convergence Zone (ITCZ), a belt with weak winds and heavy rainfall forming a barrier for eolian dust between southern and northern sources. As southeastern trade winds blow across the equator, the change in direction of the Coriolis Effect causes divergence along the equator that results in a depression in surface

69 topography and creates the pressure gradient that together produce a geostrophic flow, i.e. the South Equatorial Current (SEC). In the Pacific, the waters flowing in the SEC originate from the Peru-Chile Current (PCC) and Equatorial Undercurrent (EUC). The strength of the SEC mimics the strength of the southeastern trade winds, while the North Equatorial Current and North Equatorial Countercurrent change intensity in response to the position of northeastern trade winds (Wyrtki, 1974). The sediments of the eastern equatorial Pacific Ocean are known to be sensitive recorders of oceanographic changes and record a complex interplay of ocean chemistry, productivity, climate, and plate tectonics (van Andel et al., 1975a, b). The high sedimentation rates are engendered by an upwelling-driven, high biological productivity (van Andel et al., 1975a, b), which is estimated to contribute 18-56% to the global new production (Chavez and Barber, 1987). This upwelling, resulting from the equatorial divergence under the influence of the Coriolis Effect, provides major nutrients for phytoplankton, the primary producers in ocean surface waters and major contributors to deep-sea sediments. The bulk nutrients, however, are brought to the equator by the PCC with a shallow-water source (~50 m; Wyrtki, 1981) of a subantarctic origin. Thus the primary productivity in this region depends on the intensity of upwelling and the supply of nutrients imported at thermocline depths. The latter, in turn, is controlled by the wind stress along the equator that affects the depth of the thermocline. Nutrients in this region, however, are not depleted by phytoplankton, a condition known as “high nutrient, low chlorophyll” (HNLC) (Minas et al., 1986), which is believed to be a result of iron limitation in these waters (Martin, 1990; Martin et al., 1991). Iron is essential for all life although in small amounts (Weinberg, 1989). This need in open oceans can usually be met by the fallout of Fe-rich eolian dust (Martin and Gordon, 1988), although iron concentrations are very low due to its insolubility in oxygenated seawater, hence its fragile and transient bioavailability in the marine ecosystem.

Samples, Methods, and Age Constraints Samples were taken aboard JOIDES Resolution at a spacing of 2/section (~ 75 cm) for all cores from Hole 1256B. Stable oxygen- and carbon-isotopic analyses were carried out on each sample above 150.50 meters below sea floor (mbsf). Below this depth, one sample per 4-10 m was selected for isotopic analysis to illustrate isotopic fluctuations prior to the carbonate crash.

70 Samples for oxygen- and carbon-isotope study were prepared on bulk sediments by grinding dried sediment to a homogenous powder and baking at 425°C for 2 hours in vacuum. The stable carbon- and oxygen-isotopic ratios of the baked samples were analyzed using a Gas Bench II Auto-Carbonate device interfaced to a Finnigan Mat Delta plus XP mass-spectrometer. All isotope results were calibrated against international and internal laboratory standards, and expressed in standard delta notation relative to the Peedee belemnite (PDB) standard. Analytical precision for isotope analyses was better than ±0.1 per mil. The carbonate concentration in each sample was derived by calibrating signal amplitude of the first CO2 peak against those of carbonate standards, which are assumed to be pure carbonate. CaCO3 MARs 2 (g/cm k.y.) were calculated in using the following equation: CaCO3 MAR = (CaCO3 % x Bulk density x Linear sedimentation rate). The age model by Jiang and Wise (2007) was followed in this study, which integrated paleomagnetic data from the upper 100 m and biostratigraphic datums from the entire sequence by best linear fit. Assuming a constant sedimentation rate, age of an individual sample was obtained by extrapolating the sedimentation rate in the specific interval to the depth of this sample.

Results: Variations in Bulk Isotopes and Carbonate Accumulation Rates

A bulk-sediment stable-isotope record spanning the entire sedimentary sequence is

presented with CaCO3 MARs in Figure 2A. The oxygen-isotope record shows marked variations, ranging from -2.7‰ to 0.6‰ (Fig. 3-2). The fluctuations of δ18O values show three major patterns: (1) an irregular but consistent long- term increase from -2.2‰ to 0.1‰ prior to 11.5 Ma, (2) a sudden decrease starting at ~11.5 Ma followed by two stepwise decreases reaching a nadir at ~10.0 Ma, and then (3) a swing back and forth thereafter with a long-term increase. Significant fluctuations are also present in the carbon-isotope record, with a range from - 1.4‰ to 2.8‰ (Fig. 3-2). δ13C values are high and increase upsection just prior to 11.3 Ma, decrease sharply thereafter, recover somewhat from 11.1-11.7 Ma, decrease sharply again and remain low to the lowest point at ~8.5 Ma, then recover again thereafter. Two major negative excursions can be recognized at ~11.3 Ma and ~10.7 Ma with a net reduction in δ13C values over

71 a range of ~1.7‰ and ~1.6‰, respectively. Carbon-isotope values show little correlation with oxygen isotope records in the studied interval (Fig. 3-2B). 2 CaCO3 MARs range from 0 to 7.06 g/cm k.y. Variations in CaCO3 MARs follow the trends in stable carbon isotopes prior to 5 Ma (correlation coefficient R = 0.87) , but this positive correlation almost disappears thereafter (R = 0.42; Fig. 3-2B).

Discussion

Bulk-carbonate Isotopes Stable oxygen- and carbon-isotopic measurements of single foraminiferal species have long been used in paleoceanographic and paleoclimatic studies due to the simpler interpretation of these data. In contrast, bulk-carbonate isotopic ratios reflect weighed-average signals of different source materials, which complicates the interpretation and compromises the application of these data. Under certain circumstances, however, when the geological and sedimentary settings rule out enough such complications, bulk-carbonate isotopic records can faithfully reproduce trends from a single foraminiferal species (Shackleton et al., 1993; Schrag et al., 1995); bulk-carbonate has also been used for stable-isotope analysis. Moreover, bulk-isotopic analysis becomes the sole resort where the samples are too well-lithified to allow separation of foraminifers, where foraminifers are sparse (which is the case here), and/or where very high sample resolution is desired over long intervals. Site 1256 on the Cocos Plate is separated from the west coast of Central America by the Middle American Trench, which traps most terrestrial sediments shed from the continent. The sedimentary sources, therefore, are mainly biogenic calcite and silica as indicated by the relationship of MARs between bulk sediments and the main sedimentary components (Fig. 3-3). These biogenic materials are produced almost entirely in surface waters, and the benthic component is minor (van Andel et al., 1975a, b). Furthermore, smear-slide examinations have shown that the sediments contain highly abundant calcareous nannofossils, abundant siliceous microfossils, and extremely rare foraminifers (Shipboard Scientific Party, 2003c; Jiang and Wise, this volume). The survival and preservation of calcareous nannofossils are likely linked to the faecal-pellet transport through the water column, where organic coatings mitigate coccolith dissolution in the water column and at the sea floor (Honjo, 1976). Thus, calcareous

72 nannofossils contributed the great majority of the carbonate and consequent stable-isotopic signals. The influence of different vital effects of calcareous nannofossil species is expected to be minor. Stoll (2005) has shown that nearly mono-generic nannofossil isotopic records closely parallel those from bulk carbonate, that the inter-specific vital effect in nannofossil isotopic measurements is small, and that variable nannofossil assemblages do not significantly bias the isotopic records. Schrag et al. (1995) further demonstrated that bulk isotope data are less sensitive to compositional variations. In such a case where the nannofossil assemblage is dominated by reticulofenestrids before, during, and shortly after the carbonate crash (Jiang and Wise, this volume), bulk-carbonate isotopes can be reliably used to investigate the cause of this event.

Carbonate Diagenesis The dominant lithology in Hole 1256B is unconsolidated calcareous-nannofossil ooze (Shipboard Scientific Party, 2003c), consisting mainly of nannofossil skeletons of low-Mg calcium carbonate resistant to dissolution. Foraminifers are very rare, insufficient for isotopic analysis. Low-Mg carbonate is thermodynamically stable in deep, cold seawater and resistant to diagenetic alteration after deposition (Schlanger and Douglas, 1974); therefore, pelagic oozes remain virtually unlithified until buried to a certain depth at which point CaCO3 is released to the pore waters and re-precipitates as intra-particle fill, exterior overgrowth, and inter-granular cement (Garrison, 1981). Diagenetic alterations, including dissolution and postdepositional diagenesis, have potential for altering isotopic signals originally preserved in carbonates. Dissolution preferentially removes 13C, resulting in a depletion of 0.2‰ as observed in benthic foraminifers (McCorkle et al., 1995), though theoretically higher depletion may occur under Rayleigh distillation. Burial diagenesis, increasing with increasing CaCO3 content and burial depth, and diagenetic alteration by underlying basalts tend to progressively deplete heavier 18O (Frank and Bernet, 2000; Schrag et al., 1992). Thus, severe diagenetic alteration produces isotopically lighter calcite with well- coupled stable oxygen and carbon values. Carbon-isotope data likely have undergone little diagenetic alteration based on the following observations. (1) The calcareous nannofossil assemblage is moderately to well preserved (Jiang

73 and Wise, 2007), This is especially true for the interval where the carbonate crash was documented. (2) The correlation coefficient is low between carbon and oxygen isotopes (Fig. 3-2, B1), which otherwise indicates a precipitation of secondary isotopically lighter calcite cements attributed to an early diagenetic overprint (Jenkyns, 1974) or deep burial (Jenkyns and Clayton, 1986; Jenkyns, 1995). (3) An inverse relationship is absent between organic carbon and carbon- isotope data (Figs. 2 and 3; Jenkyns and Clayton, 1986). Therefore, carbon-isotope values are considered to represent a primary environmental signal. Oxygen-isotope values are prone to diagenetic alteration during burial diagenesis because oxygen isotopes show a significant temperature-dependent fractionation (Anderson and Arthur, 1983; Marshall, 1992). However, Schrag et al. (1995) has demonstrated that the effect of rapid calcite precipitation is small for the biogenic carbonates from mid-latitude Atlantic, because primary oxygen-isotope values in carbonates are close to isotopic equilibrium with cold pore fluids. As today, Site 1256 was located during the middle Miocene beneath the equatorial divergence and under a strong influence of upwelling of cold deep water, so the SSTs are considered to be comparable to those of the temperate Atlantic as indicated by today’s ocean SST pattern. The sediments at this site are shallowly-buried (250.7 m), and hydrothermal circulation is no longer a major mechanism of heat transport (Shipboard Scientific Party, 2003c). Therefore, oxygen-isotope data mainly reflect environmental signals.

Surface-water Carbonate-production Proxies Paleoproductivity reconstruction is a major area of paleoceanographic research because it places important constraints on past ocean circulation, nutrient distribution, and oceanic carbon- cycle history. Important proxies have been developed for past productivity and interpreted in terms of organic-carbon export. These include direct measurements of organic carbon in sediments, biogenic opal and calcium carbonate accumulation, foraminiferal assemblage data, geochemical tracers, and δ13C fluctuations (e.g., Müller and Suess, 1979; Berger et al., 1989; Herguera and Berger, 1991; Paytan et al., 1996; Tappan, 1968). Processes that influence these estimates, such as changes in seawater chemistry, dissolution, diagenesis, and nutrient availability, complicate the interpretations and thereby limit the application of these proxies. Because organic carbon itself is subject to influence by downslope transport and degradation

74 13 (Fig. 3-3), CaCO3 MARs and δ C values have been employed in a multi-proxy strategy in this study to assess changes in paleoproductivity. As the major components in the sediments are biogenic skeletons of organisms dwelling in the uppermost water column (Fig. 3-3), the geochemical composition of the sediments

predominantly reflects a surface-water signal. Thus, CaCO3 MARs reflect carbonate production by surface-dwelling, calcite-secreting organisms when dissolution is a minor effect. Interpretation of δ13C, however, depends on the specific geological and oceanographic settings. The bulk carbonate in this study was produced nearly exclusively in the surface water; therefore, variations in the bulk δ13C values should reflect fluctuations therein, which are the product of the interplay between phytoplankton photosynthesis and surface-water chemistry as influenced by seawater alkalinity and the dissolved-carbon 13C/12C ratio. Phytoplankton dwell in the oceanic euphotic layer and preferentially take up 12C during photosynthesis, producing organic matter with δ13C values of -20~-23‰ and a relatively 13C-enriched dissolved inorganic- carbon pool in surface waters. Seawater pH influences the concentrations of different forms of 13 2- dissolved inorganic carbon, the δ C values of which decrease with increasing seawater [CO3 ] (Spero et al., 1997). 13 In the present study, the close coupling between the CaCO3 MARs and δ C values prior to 5 Ma and decoupling thereafter (Fig. 3-2B) reflect a switch of dominance between carbonate production and dissolution. Prior to 5 Ma, carbonate production controls the sedimentary patterns at Site 1256, and δ13C variations predominantly reflect changes in the standing stock of carbonate-producing organisms in the surface waters (Figs. 2 and 3). After 5 Ma, dissolution dominates this site, probably resulting from the effective blocking of the Panama Seaway since then (Haug and Tiedemann, 1998). The large excursions observed in δ13C values could not have arisen from enhanced 2- dissolution. Dissolution occurs when there is a reduction in [CO3 ] in the surface water, the water column that the carbonates travel through, and the bottom water. Changes in seawater 2- 13 [CO3 ] in intermediate and deep waters, however, do not significantly alter the δ C signal of carbonates produced, as observed in foraminifers (McCorkle et al., 1995). If such reduction in 2- 13 18 [CO3 ] occurred in surface waters, it would have greatly increased the δ C and δ O values of carbonate (Spero et al., 1997), a situation that is opposite the observations here (Fig. 3-2).

75 Therefore, no matter where severe dissolution occurs, it could not produce the negative excursions in δ13C and δ18O values observed in this study. δ13C values faithfully represent variations in the surface-water standing stock of calcite makers.

Carbonate Crash Timing of Carbonate Crash

CaCO3 sedimentation at Site 1256 shows several extreme lows between 12 Ma and 8 Ma, which were initiated at ~11.3 and 10.6 Ma, geologically synchronous with the onset of the carbon-isotope excursions (Fig. 3-2). The deepest drop is at 9.6 Ma with carbonate accumulation virtually ceasing. This is temporally comparable with the “nadir” of the carbonate crash at other ODP/Deep Sea Drilling Program (DSDP) sites in this region (Lyle et al., 1995; Farrell et al., 1995), whereas the five carbonate minima in the Caribbean occurred ~1 Ma earlier (12-10 Ma; Roth et al., 2000). These phenomena were previously ascribed to enhanced carbonate dissolution as a consequence of changes in deep-water circulation (e.g. Lyle et al., 1995; Roth et al., 2000).

Carbonate Accumulation in Open Oceans Accumulation rates of biogenic carbonate depend on surface-water productivity, dissolution in the water column and, after they reach the bottom, dilution by non-carbonate components. The production of biogenic carbonate in the open ocean relies on the abundance and distribution of calcite-precipitating organisms (e.g. coccolithophore and foraminifera) which, in turn, depends in large part upon nutrient supply and temperature in their habitats. A high Mg/Ca ratio and low absolute concentration of Ca also limit population growth of coccolithophores (Stanley et al., 2005). On the ocean floor, dissolution of calcium carbonate in seawater is determined by seawater pH values influenced by temperature, pressure and the partial pressure of CO2. Only above the carbonate compensation depth (CCD) can carbonate accumulate. The depth of the CCD varies in different ocean basins depending on bottom-water chemistry and carbonate supply from surface water (Wise, 2003). The former explains a present-day CCD in the Pacific shallower than in the Atlantic, the latter a shallower CCD in high latitudes where carbonate production is low. High biological productivity tends to stimulate higher rates of carbonate production and to depress the CCD. However, if high productivity is engendered by extremely high concentrations of nutrients,

76 a condition that favors diatoms and dinoflagellates, the CCD tends to shoal due to addition of

CO2 to the bottom water and/or acid production resulting from the degradation of organic matter (Emerson and Bender, 1981; Archer, 1991a, 1991b). This scenario may occur in response to changes in climate (Dymond and Lyle, 1985). Carbonate in sediments can be diluted by increases in non-carbonate components, which causes the decrease in carbonate weight content. This generally occurs in areas that are in close approximation to terrigenous delivery, such as river mouths, coastal waters, and many passive- margin continental shelves. Dilution of carbonate can also occur in the open ocean where siliceous organisms thrive under condition of high nutrient supply.

Previous Models for the Carbonate Crash Multiple causative mechanisms have been proposed for the carbonate crash. Severe dissolution in the equatorial Pacific and Caribbean has been attributed to changes in bottom- water chemistry. The underlying hypotheses are either (1) the early phase of constriction of the 2- Panama Seaway that restricted the exchange of high-[CO3 ] deep waters between the Pacific and the Caribbean (Lyle et al., 1995; Farrell et al., 1995), or (2) the intensified influx of corrosive water in response to strengthened global thermohaline circulation linked to enhanced production of NADW (Roth et al., 2000). Although the “nadir” of the carbonate crash occurred ~1 Ma earlier in the Caribbean than elsewhere, the comparable nature and time-overlap of its occurrences suggest a common cause associated with changing oceanic circulation (Roth et al., 2000). This does not exclude other competitive mechanisms, such as “dilution” by terrigenous sediments (Diester-Haass et al., 2004) or an opal component (Westerhold et al., 2003; Böhm and Dullo, 2000) for those “carbonate crashes” reported elsewhere (i.e. the southern and eastern South Atlantic, southern Indian Ocean). Phytoplankton-community restructuring against calcite-producing organisms could also produce severe reduction in CaCO3 MARs, even though the overall surface-water productivity remains constant (Dymond and Lyle, 1985). At Site 1256, the negative excursions in δ13C and δ18O excursions coincide prior to the carbonate crash. At this time, an oxygen-isotope signal is supposed to record a combination of surface temperature, ice volume, and seawater chemistry, the latter two of which tend to enrich heavier isotopes in response to the development of the East Antarctic Ice Sheet (EAIS) (Zachos et al., 2003) and, if any, the dissolution-related decrease in

77 2- 13 seawater [CO3 ]. Thus, δ C excursions observed at Site 1256 likely represent sudden SST increases, a condition that favors warm-water specialists/calcite-producing organisms (McIntyre and Bé, 1967; 1970). Similar observations, especially the coeval excursion in δ13C and δ18O values at 11.3 Ma, have also been documented in other nearby ODP/DSDP sites (Shackleton and Hall, 1984 and 1995) and in the Caribbean (Mutti, 2000). In other words, these δ13C excursions represent a sharp decrease in standing stock of calcareous phytoplankton in surface waters. Evidence for increased dissolution across the carbonate crash includes decreased coarse calcareous fraction, increased benthic/planktonic foraminifer ratios, and deteriorated preservation of calcareous fossils. Unfortunately, these are indirect proxies and should be cautiously placed under a sedimentary and geochemical context. The drop in the sand-sized fraction recorded at Sites 998 and 999 was attributed to enhanced fragmentation of foraminiferal tests in more corrosive water columns (Roth et al., 2000), while the same phenomenon observed in southeast Atlantic (Sites 1085 and 1087) was related to sea-level regression and consequent increased terrestrial input (Diester-Haass et al., 2004). The increased corrosiveness of the water column itself, however, does not necessarily demand an influx of corrosive waters, because a reduction in carbonate production in the surface water would have the same effect. The increase in the ratios of benthic to planktonic foraminifera in southwest Atlantic seems to be controlled more by nutrient condition than dissolution (Diester-Haass et al., 2004). At Site 1256, the calcareous nannofossil assemblages exhibit their best preservation just prior to the carbonate crash nadir (Jiang and Wise, this volume). This does not necessarily mean no dissolution occurred during the two dramatic drops in carbonate MARs, because the preservation is closely associated with the presence and abundance of diatoms, which release silica to pore waters and thus inhibit dissolution and/or precipitation of calcite during diagenesis (Wise, 1977). The constriction of the Panama Seaway in the late middle Miocene limited communication between the Atlantic and Pacific at intermediate- and deep-water levels (Duque-Caro, 1990), and could have caused basin-to-basin isotopic fractionation (Lyle et al., 1995; Roth et al., 2000), a scenario seen today. The modern Atlantic is filled with young, well-oxygenated water due to the production of NADW, whereas the Pacific has older water originating from the North Atlantic and southern high latitudes. As these waters age, the organic matter therein decomposes and releases 12C-depleted carbon into seawater, which is responsible for the more negative δ13C

78 values in eastern Pacific relative to the western Atlantic (Fig. 3-4; Kroopnick, 1985). However, the following perceptions raise several questions regarding this mechanism. It takes about 1500 years for deep water sinking in the North Atlantic to reach the north Pacific (Manighetti, 2001), which cannot explain the 1-million-year lead for the carbonate crash in the Caribbean. The water brought up to the surface by equatorial divergence-driven upwelling has a shallow-water source above the thermocline with a mean depth of 50 m (Vossepoel et al., 1999). In fact, no matter where the ultimate source of these waters is, the North Atlantic or circumpolar Antarctic, the δ13C difference between these sources today and the eastern Pacific is no more than 0.7‰ (Fig. 3-4). This alone could not have caused the δ13C excursions of >1.6‰ accompanying the carbonate crash documented at Site 1256. The carbonate crash was observed mostly from records retrieved from >3000 m water depth, a depth level that is highly sensitive to fluctuations in the CCD in the eastern equatorial basins (Lyle et al., 1995). The previously proposed mechanisms can hardly account for its occurrences at very shallow sites (e.g. Sites 1000 and 1241) in water depths well above the modern CCD, hence the least susceptible to the influence of shoaling CCDs. As a matter of fact, carbonate production played a dominant role over dissolution as 13 evidenced by the close parallelism in CaCO3 MARs and δ C values in the 14-5 Ma interval and by the absence of deterioration in preservation of calcareous nannofossils during the crash (Raffi and Flores, 1995; Jiang and Wise, this volume). In other words, the crash was not a dissolution event, but an low productivity event. This conclusion is in line with the lack of basin-wide occurrences of the crash in the Pacific (Lyle, 2003).

Proposed Model for the Carbonate Crash Based on these observations and discussions above, we speculate that a reduction in the standing stock of carbonate-producing organisms, likely induced by a reduction in nutrient availability, triggered the worldwide carbonate crash at the middle/late Miocene boundary. The model proposed here emphasizes that carbonate supply to intermediate and deep waters is primarily a function of changes in carbonate production, and as a consequence, affects the corrosiveness of these waters on carbonate and preservation of carbonate. The change of carbonate production is associated with changing global surface-water circulation, which, in turn, is induced by the seaway-controlled inter-ocean exchange of water masses.

79 During late Cenozoic, prior to the segmentation of the once virtually continuous circum- equatorial current, nutrient-rich intermediate water was tapped out to the surface where it sustained high biological productivity along equator. This flow began to fragment due to the uplift of the Panama Isthmus at 12.9-11.8 Ma (Duque-Caro, 1990) and the closure of the Indonesian Seaway at 12-11 Ma (Keller, 1985; Romine and Lombari, 1985; Kennett et al., 1985). The latter occurred when subduction developed on the east and west sides of New Guinea (Hall, 2001). The difference in timing can explain the ~1-million-year lead of the carbonate crash in the Caribbean relative to the Pacific and Indian Oceans. Though the Indonesian Seaway was effectively blocked at about 17-15 Ma, Indo-Pacific communication was still possible through small passages during times of high sea level (Nishimura and Suparka, 1997). This transport of surface water, termed as the Indonesian Throughflow (ITF) today, carries warm and fresh tropical-Pacific surface water through this passage into the Indian Ocean and creates a warm pool in the western Indian Ocean. The global sea-level drop at the middle/late Miocene boundary (Haq et al., 1987) completely blocked this Pacific-to-Indian surface-water transport, switching the locale for warm-water accumulation to the western Pacific, creating a greater warm pool therein. The eastward spread of the warm-pool water strengthened the EUC system (Kennett et al., 1985), resulting in warm SSTs in the central and eastern Pacific and reduced equatorial upwelling of colder subsurface water, both of which contribute to the creation of a δ18O excursion. The switch of dominance to nutrient-poor warm surface water caused a sudden reduction in biological productivity, which is represented by a δ13C excursion. This event is analogous to today’s El Niño. A similar scenario may have occurred in the Caribbean and Atlantic when the Panama Seaway suddenly became restricted at 12.9-11.8 Ma. Because blockage of the ITF cut off Pacific-to-Indian Ocean heat transport, this should have triggered an overall warming in the tropical Pacific and cooling in the southern Indian Ocean, respectively. This is evidenced by an abrupt disappearance of foraminiferal and radiolarian provincialism across the equatorial Pacific (Keller, 1985; Romine and Lombari, 1985; Kennett et al., 1985), and a positive δ18O excursion in shallow-dwelling planktonic foraminifers at DSDP Sites 216 and 237 (Vincent et al., 1985). A temporally similar warming of the entire water column was recorded at DSDP Site 289 (Gasperi and Kennett, 1993). The eastern spread of warm surface water warmed not only the east equatorial Pacific, but also southeast Pacific off

80 Chile (Tsuchi, 1997). These perceptions are consistent with the results of a near-global ocean general circulation model for the circulation and thermal structure of the Pacific and Indian Oceans with open and closed Indonesian passages from 1981 to 1997 proposed by Lee et al. (2001). It is worth mentioning that the evolution of global climate entered a full icehouse mode during the Neogene (Zachos et al., 2003). Cold climate narrows carbonate producing areas, rendering carbonate production in tropical oceans more important; small changes in production therein would greatly affect total carbonate supply to deep waters. After the full development of the EAIS by the middle Miocene (Shackleton and Kennett, 1975; Kennett et al., 1985; Woodruff and Savin, 1991; Zachos et al., 2001) (North Hemisphere glaciations did not occur until ~7 Ma [Fronval and Jansen, 1996]), the southern trades winds began to prevail over their northern counterpart, causing a northward shift of the ITCZ, an analogue of the initiation of modern North Hemisphere summers. The intensification of southeastern trade winds during the Neogene is consistent with the coarsening of eolian grain size recorded from the subtropical south Pacific (Rea and Bloomstine, 1986). Positions farther north than present (5°N) of 10-12°N, ~22-24°N, and 13-14°N have been suggested for the ITCZ around the middle-late Miocene boundary by Flöhn (1981), Rea (1994), and Shipboard Scientific Party (2002), respectively. This northward shift of the ITCZ can potentially wash out eolian dust at the northern mid-latitudes, mitigating the iron-limited condition and stimulating biological production. This scenario is consistent with the records from North Pacific (Snoeckx et al., 1995). The eastern equatorial Pacific is presently an HNLC region because of iron limitation on phytoplankton (Martin et al., 1991). This region did not suffer from iron deficiency during the early and middle Miocene when there was intense activity in circum-Caribbean and Central American volcanism as evidenced by tephra accumulation rates (Sigurdsson et al., 2000), but it has since then. Once exposed to seawater, volcanic ash can fertilize the open ocean by releasing large amounts of macronutrients and bioactive trace metals (Frogner et al., 2001). According to our model, with the closure of the Indonesian Seaway by the sea-level drop at the middle/late Miocene boundary, warm surface water began to pile up in the western Pacific. The formation of a larger warm pool strengthened the EUC system. The eastward spread this water triggered a prominent “El Niño”, reducing upwelling and then nutrient availability. At the same time, prevailing southeastern trade winds across the equator deflected the delivery of

81 volcanic ash from Central America and the circum-Caribbean, where volcanism was coincidently reduced sharply (Sigurdsson et al., 2000). The decrease in availability of macronutrients and micronutrients resulted in a drastic reduction in primary productivity in the eastern and central equatorial Pacific as well as the Caribbean, as evidenced by the carbon-isotope excursions at Site 1256 and the synchronous variations in MARs of carbonate and volcanic ash throughout the Neogene at Sites 998 and 999 (Fig. 3-5). The CCD in the Neogene Pacific is a product of the balance between regional production and basin-wide dissolution (Lyle, 2003). The modern deep Pacific below 1500 m is essentially one water mass (Joyce et al., 1986; Talley and Roemmich, 1991); hence the entire Pacific Ocean floor is bathed by the same water mass and carbonate dissolves primarily at the ocean floor (Edmond, 1974; Walsh et al., 1988). On the one hand, dissolution rates should be similar everywhere and there should be a base-level CCD for all regions in the Pacific. The variable depth of the CCD in the Pacific, however, originates from biogeographic differences (Lyle, 2003). That is, although there is a base level, the CCD may be depressed or elevated depending on surface productivity. On the other hand, although there has always been during the Neogene an Antarctic deep-water source into the Pacific basin, no convincing evidence exists for stronger flow at the middle/late Miocene boundary (Lyle et al., 1995). The consequent near-constant deep-water chemistry, plus the close correlation between CaCO3 MARs and biological productivity, suggests that the carbonate crash in eastern equatorial Pacific and likely in other regions is not a dissolution event, but one of low productivity.

Conclusions

The late/middle Miocene carbonate crash has previously been considered only a dissolution event associated with the production of the North Atlantic Bottom Water and ventilation via the 13 Panama Seaway. A coupling between δ C and CaCO3 MARs data in the 5-14 Ma interval (correlation coefficient = 0.87) observed at ODP Site 1256 suggests a dominant role for carbonate production by calcite-secreting organisms in surface waters, while the decoupling thereafter probably indicates an enhanced dissolution caused by the effective closure of the Panama Seaway. Therefore, the coeval occurrences of negative δ13C excursions and stages of

82 sharp reduction in CaCO3 MARs during the middle/late Miocene carbonate crash, points to a causative mechanism related to surface-circulation-induced low productivity. We speculate that the major middle/late Miocene sea-level drop caused the complete closure of the Indonesian Seaway. The blockage of the ITF resulted in a piling-up of surface warm water in the west Pacific, strengthening the EUC system. We further speculate that the eastward spread of this nutrient-poor water warmed the SST and reduced upwelling in the central and eastern Pacific, primarily triggering a reduction in the standing stock in calcareous phytoplankton. The synchronous reduction in Central America and circum-Caribbean volcanism and deflected delivery of volcanic ash would have further deprived these regions of trace elements, which added to lowered surface-water carbonate production. The surface-water warming and reduced upwelling is documented by a negative excursion in δ18O values. The reduction in carbonate supply to the deep waters caused a rapid shoaling of the CCD and triggered the carbonate crash.

The close coupling between CaCO3 MARs and biological productivity suggests that the carbonate crash is not a dissolution event, but one of low productivity. In order to test the mechanism proposed here, further work needed includes comparisons of the temperature variations at sea surface, thermocline, and if possible, bottom. This is possible by comparing stable-isotope records spanning the carbonate crash from shallower localities in east and west Pacific, respectively, where foraminifera are supposed to be well preserved. To simplify the interpretation of isotopic data, it is ideal to analyze single foraminiferal species from niches in different water depths.

Acknowledgements

We wish to thank all the participants of ODP Leg 206 for their contributions relevant to this paper, and Dana Biasatti from the Stable Isotope Lab in National High Magnetic Field Lab for her helpful assistance in running samples. We thank Drs. Gerald J. Dickens, Mitchell W. Lyle, Larry C. Peterson … for helpful and constructive reviews. This research used samples and/or data provided by ODP. ODP is sponsored by the U.S. National Science Foundation (NSF) and participating countries under management of Joint Oceanographic Institutions (JOI), Inc. Funding for this research was provided by a USSAC grant to S.J. and S.W.W. General lab support was provided by NSF-OPP 0126218 to S.W.W.

Figure 3-1. Locations of Site 1256 and other ODP/DSDP sites discussed in this study and sketch of major world surface currents. Equatorial Undercurrent in the Pacific Ocean is also shown as a series of dashed arrows. Gray lines across the map show the annual shift of the Intertropical Convergence Zone: Solid = July, Dash = January. Gray arrow = Equatorial Countercurrent.

Figure 3-2. Relationship between carbonate mass accumulation rates (g/cm2 k.y.) and stable oxygen and carbon isotopes.

Figure 3-3. Relationship of mass accumulation rates between bulk sediments and the main sedimentary components (from Jiang and Wise, 2007).

Figure 3-4. Correlation between CaCO3 and volcanic ash MARs from ODP Leg 165 (data from Peters et al., 2000).

Figure 3-5. δ13C of the modern western Atlantic (A), eastern Pacific (B), and circumpolar Antarctic (C) (from Kroopnick, 1985).

CHAPTER 4

SURFACE-WATER CHEMISTRY AND FERTILITY VARIATIONS IN THE TROPICAL ATLANTIC ACROSS THE PALEOCENE/EOCENE THERMAL MAXIMUM AS EVIDENCED BY CALCAREOUS NANNOPLANKTON FROM ODP LEG 207, HOLE 1259B

Abstract

Calcareous nannoplankton assemblages at Ocean Drilling Program Site 1259 at mid-bathyal depths on Demerara Rise (western equatorial Atlantic) underwent an abrupt and fundamental turnover across the Paleocene/Eocene Thermal Maximum (PETM) ~55.5 m.y. ago. The P/E boundary event is marked by a dissolution interval barren or nearly-barren of nannofossils due to the rapid acidification of the world oceans. Toweius, Fasciculithus, and Chiasmolithus sharply decrease at the onset, whereas Chiasmolithus, Markalius cf. M. apertus, and Neochiasmolithus thrive immediately after the event, which also signals the successive first appearances of Discoaster araneus, Rhomboaster, and Tribrachiatus. The environmental indications of these changes were further investigated by correspondence analysis on quantitative nannofossil counts based on two main factors extracted.

The PETM event has been attributed to CO2-forced greenhouse effects. At Site 1259, the elevated pCO2 and subsequent lowered surface-water pH values at the onset of the PETM caused intensive carbonate dissolution, producing the nannofossil-barren interval. The chemically stressed habitats may well have also induced the evolution of ephemeral nannofossil “excursion

89 taxa”, such as Rhomboaster and malformed discoasters (Discoaster araneus and Discoaster anartios). Based on its sudden increase, Markalius cf. M. apertus is considered to have been a local opportunistic species that took advantage of the surface-water changes. At the same time, a presumably higher runoff from continental areas fertilized the western equatorial Atlantic as indicated by an increase in the abundance of r-mode specialists preferring high-nutrient conditions, such as Chiasmolithus, Coccolithus pelagicus and Hornibrookina arca. Contrasts between the results of this study and previous work at ODP Site 690 in the Southern Ocean, the New Jersey continental margin, and the central paleoequatorial Pacific further demonstrate that the response to the PETM can be influenced by local differences in geologic setting and oceanographic conditions. Key Words: calcareous nannofossils; paleoecology; Paleocene/Eocene thermal maximum; correspondence analysis; surface-water chemistry; paleoproductivity

Introduction

The Paleocene/Eocene Thermal Maximum (PETM; formerly termed the Late Paleocene Thermal Maximum, or LPTM by Zachos et al., 1993) was a catastrophic, transient (<220 kyr [Norris and Röhl, 1999; Röhl et al., 2000; Farley and Eltgroth, 2003]) climate event ~55.5 m.y. ago. This global warming elevated sea-surface temperature (SST) by 5ºC in the tropics and as much as 9ºC at the high latitudes (Kennett and Scott, 1991; Zachos et al., 2003), whereas bottom water warmed by 4-5ºC (Thomas and Shackleton, 1996; Thomas et al., 2000). This event has been widely accepted as one of the best examples of deep-time, rapid greenhouse-forced warming. The most compelling evidence for this extreme event is a coeval carbon-isotope excursion (CIE) of ~3‰ recorded in global marine and terrestrial systems, suggesting a massive perturbation to the global carbon cycle (e.g., Kennett and Stott, 1991; Koch et al., 1992; Thomas and Shackleton, 1996; Beerling and Jolley, 1998; Bains et al., 1999; Norris and Röhl, 1999). The large magnitude and pattern of rapid decrease and gradual recovery through the CIE indicates multi-pulse injections of isotopically depleted carbon into the ocean and atmosphere (Dickens et al., 1997; Bains et al., 1999; Dickens, 2000). The most plausible source of this carbon is the sudden dissociation of methane hydrates from continental shelves and slopes (Dickens et al.,

90 1995, 1997; Katz et al., 1999), whereas thermogenic methane from hydrothermal vents (Svensen et al., 2004) and carbon dioxide from global wildfires (Kurtz et al., 2003) or an extraterrestrial comet (Kent et al., 2003) have also been suggested. Methane, or its subsequent product of oxidation, carbon dioxide, would have immediately contributed to greenhouse warming and changes in ocean chemistry. The abrupt climatic change associated with the PETM triggered massive turnovers in oceanic benthic and planktonic organisms as well as in terrestrial vegetation and mammals, and caused a widespread rapid shoaling of the calcite compensation depth (CCD) (Zachos et al., 2005). Deep-water warming and subsequent oxygen deficiency may have led to the extinction of 30%–50% of benthic foraminiferal species (Tjalsma and Lohmann, 1983; Thomas, 1990; Kennett and Scott, 1991; Thomas and Shackleton, 1996), the rising temperature to the evolutionary radiation of land mammals and vegetation (Koch et al., 1992; Clyde and Gingerich, 1998; Meng and McKenna, 1998; Wing and Harrington, 2001; Jaramillo, 2002; Gingerich, 2003) and diversification of planktonic foraminifers (Kelly et al., 1996, 1998; Berggren and Ouda, 2003), and the changes in oceanic surface-water chemistry to the evolution of ephemeral “excursion taxa” among planktonic foraminifera and nannofossils (Bybell and Self-Trail, 1995; Kelly et al., 1996; Kahn and Aubry, 2004). The responses and/or strategies of surface-dwelling and benthic organisms to the PETM environmental changes appear to have been fundamentally different. Because of very limited high-resolution investigations using calcareous nannofossils (e.g. Angori and Monechi, 1996; Bralower, 2002; Tremolada and Bralower, 2004; Raffi et al., 2005; Gibbs et al., 2006) and their Sr/Ca ratios (Stoll and Bains, 2003), and the lack of consensus on productivity variation during the PETM from different proxies even at the same site (Bains et al., 2000; Bralower, 2002; Stoll and Bains, 2003), just how phytoplankton, the primary producers in ocean surface waters and major carbonate contributors to deep-sea sediments, reacted to the PETM, is still not well understood. Recent comparisons, however, suggest that their response were different in continental margin vs. open ocean settings (Gibbs et al., 2006). The equatorial Atlantic Ocean has been long known to be an important repository of information for Earth’s history. Sediments that have accumulated beneath the divergence-driven upwelling system provide the most continuous high-resolution archives and sensitive recorders of oceanographic change, which document a complex interplay of ocean chemistry, productivity,

91 climate, ocean-atmospheric circulation, and plate tectonics (van Andel et al., 1977). Because the sediments are shallowly buried, crop out on the seafloor in places, and contain microfossils generally better-preserved than elsewhere, they are ideal for studying Earth’s past oceanographic and climatological fluctuations. Ocean Drilling Program (ODP) Leg 207 was committed to recover relatively expanded, shallow-buried sediments from the tropical Atlantic (Fig. 4-1) that could be used for paleoceanographic study of Cretaceous and Paleogene global-change events (Shipboard Scientific Party, 2004a). One of the primary scientific objectives was to evaluate the biotic turnover across the PETM. The Paleocene/Eocene (P/E) boundary was recovered at all five sites, and represented by a dark green clay-rich bed in sharp contact with underlying chalk, providing an excellent opportunity to investigate this event. We present here a quantitative study of calcareous nannofossil assemblages across the PETM from ODP Site 1259 (Demerara Rise, South America; Fig. 4-1), located along the outer continental margin in a water depth of 2354 meters below sea level, in an attempt to investigate the biotic impacts of the PETM on marine phytoplankton. We then discuss the surface-water variations across the PETM in light of these data.

Materials and Methods

Samples used to study calcareous nannoplankton were taken using a U-channel through the archive half of the cores between Samples 1259B-8R-2, 1-2 cm (364.21 meters below sea floor [mbsf]) and 8R-4, 148-149 cm (368.68 mbsf), sub-sampled continuously every 1 cm. A subset of samples (48 samples) was picked for this reconnaissance study at a spacing of 4-14 cm based on sample availability and proximity to the event. Preparation of smear slides followed standard techniques using Norland-61 optical adhesive as a mounting medium; unlithified samples were left in water until completely disaggregated, and no settling was used. Calcareous nannofossils were examined using a Zeiss Axiophot II microscope under crossed-polarized light, transmitted light, and phase-contrast light at 1250x magnification and identified using standard taxonomy as described by Perch-Nielsen (1985), Bybell and Self-Trail (1995), and Wise et al. (2004, CD-ROM). A JEOL JSM 840 scanning electron microscope (SEM) was employed to take digital images for more precise species identification. The biostratigraphic

92 zonation followed schemes proposed by Martini (1971; modified in Martini and Müller [1986]) and Okada and Bukry (1980) for low latitudes, and the recognition of the P/E boundary was based on several nannofossil datums (Aubry et al., 1996; Cramer et al., 2003; and Raffi et al., 2005) as well as abundance crossovers (Bralower, 2002). A total of 60 different taxa have been identified and recorded (Table 4-1), and then grouped into 13 generic and/or paleoecological categories to enhance trends for further analysis as well as to eliminate influence of assemblage preservation when many specimens can only be identified to generic levels (Table 4-2). All nannofossil taxa considered in this paper are listed in Appendix B, where they are arranged alphabetically by generic epithets. Bibliographic references for these taxa can be found in Perch- Nielsen (1985), Bybell and Self-Trail (1995), and Bown (1998). Quantitative counts involved a total of >400 specimens except for barren or nearly-barren samples. Counts were made on all specimens larger than half the size of a complete fossil in randomly selected fields of view with sample materials evenly distributed. Counting continued in the last field of view even after the overall count reached 400. Searching for zonal markers didn’t stop until three transverses of each slide (~600 fields of view) have been browsed, and the presence, if any, of these taxa was not added to the total counts and was only represented by a slash (/). Relative abundance as the percentage of each group was calculated only for those samples with total specimen counts greater than 400. Shannon-Weiner index, a diversity measurement formulated in terms of an entropy function representing a system’s degree of indeterminacy, was calculated for each sample to demonstrate the diversity of the assemblage. Assemblage preservation was monitored using visual observations of all specimens under the light microscope, and classified as follows below with transitional conditions between them. This information was further quantified using the ratio of indetermined Toweius to species- identifiable Toweius, because the identification of Toweius species mainly depends on central structures susceptible to alteration and overgrowth. G = good (little or no evidence of dissolution and/or overgrowth; primary morphological characteristics only slightly altered; specimens are identifiable to the species level). M = moderate (specimens exhibit some etching and/or overgrowth; primary morphological characteristics sometimes altered; however, most specimens are identifiable to the species level).

93 P = poor (specimens are severely etched or exhibit overgrowth; primary morphological characteristics largely destroyed; fragmentation has occurred; specimens cannot be identified at the species and/or generic level). All environmental interpretations are based on percentage variations of the selected 13 fossil groups. In order to further reduce the dimensions of the nannofossil assemblage data matrix, correspondence analysis (CA) was carried out to extract the major information on a multivariate (multi-taxa) basis. CA assumes that species often have unimodal species response curves (ter Braak and Prentice, 1988), which is more the case with a single abundance peak for most groups in the assemblages across the PETM (Fig. 4-2). This analysis was carried out in a CA program (formulated by Dr. William Parker, unpublished) plugged in to the Matlab 7 platform.

Results: Nannofossil Turnover across the PETM

In Hole 1259B, there is a sharp lithologic contact at 106 cm in Core 8R, Section 4 (Fig. 4-3; 368.26 mbsf) between a dark-green clay layer and the underlying chalk. This marks the beginning of a dissolution event represented successively upsection by a barren interval (Samples 1259B-8R-4, 103-104 cm and 97-98 cm) overlain by a nearly-barren interval (Samples 1259B-8R-4, 90-91 cm and 83-84 cm) in terms of nannofossils. Except for four barren or nearly- barren samples from this interval, all other samples yielded abundant nannofossils. Visual examination under the light microscope indicated that nannofossil assemblages are moderately to poorly preserved, and this qualitative preservation is generally consistent with the quantitative assessment (Fig. 4-4). Preservation deteriorates close to the sharp lithologic contact, where nannofossils are sparse. Preservation and diversity covary throughout the section (Fig. 4-4), a reasonable relationship that demonstrates that more taxa could be identified in better-preserved assemblages. Specimens in the Rhomboaster-Tribrachiatus plexus suffer from moderate to heavy overgrowth. This does not hamper the identification of the boundary between Zones NP10a and NP9b defined by the first occurrences (FO) of Tribrachiatus bramlettei, which is reasonably consistent with the CP9/CP8 boundary recognized by the FO of sparse Discoaster diastypus in this hole (Table 4-1). The carbon isotope excursion is a widely-accepted criterion for defining the PETM. A series of calcareous nannofossil datums has been shown to approximate this geochemical horizon

94 (Aubry et al., 1996; Bralower, 2002; Cramer et al., 2003; Raffi et al., 2005). At Site 1259, Fasciculithus decreases in diversity and abundance below the bottom of the barren interval; the Rhomboaster–Tribrachiatus lineage and Discoaster araneus evolve, and the FO of D. diastypus occurs within the nearly-barren interval. A crossover in dominance from Fasciculithus uphole to Neochiastozygus junctus and the peak abundance of the genus Discoaster occur above the dissolution interval (Fig. 4-2; Table 4-1). These observations help pin the P/E boundary to the sharp lithologic contact between Samples 1259B-8R-4, 109-110 cm (368.29 mbsf) and 8R-4, 103-104 cm (368.23 mbsf), which is compatible with and further constrains the stable bulk- isotope data with the carbon isotope excursion observed between Samples 1259B-8R-4, 107 cm (368.27 mbsf; 2.17‰) and 8R-4, 54 cm (367.74 mbsf; -0.141‰) (R.D. Norris, pers. comm., 2006). Assemblage counts indicate profound changes in nannoplankton populations within the studied interval (Table 4-1; Fig. 4-2). Coccolithus pelagicus and Toweius are dominant elements in the nannoplankton populations throughout the section, making up ~2/3 or 66% of the population before and after the PETM. Nevertheless, the nannoplankton community underwent dramatic turnover at the onset of the event. Significant changes include the following: (1) Toweius shows a dramatic drop upsection in relative abundance and never recovers to pre-event values. The relative abundance reaches an average value less than 20% of the previous values in the three samples immediately above the P/E boundary. (2) Fasciculithus and Neochiastozygus (elliptical nannoliths with a central-area bridge or structure) show a reversal in percentages across the P/E boundary. (3) The sudden appearance and rapid increase of “excursion taxa” occurs shortly after the PETM, making up ~35% of the assemblage compared to near-zero pre- event values. These taxa include Discoaster araneus, Discoaster anartios, Rhomboaster, and arguably Markalius cf. M. apertus. Correspondence analysis of the relative abundance for the entire interval investigated extracted 13 factors (Fig. 4-5A). The cumulative percent of covariance of the first two factors reaches 84.6% (Table 4-3), which indicates that this large amount of information extracted from the original relative abundance data set can be explained by these two factors. High loadings on the first two factors demonstrate strong association and non-association among taxa. On Factor 1, excursion discoasters and Markalius cf. M. apertus bear a high positive loading, Rhomboaster an intermediate positive loading, and Sphenolithus a low positive loading, whereas Fasciculithus

95 and Toweius show a high negative loading (Table 4-3, Fig. 4-5B); other groups show extremely low loading on this factor. Factor 2 shows a very high positive loading for Fasciculithus, an intermediate positive loading for excursion discoasters, and an intermediate negative loading for Coccolithus pelagicus and Hornibrookina arca (Table 4-3, Fig. 4-5B); other groups bear low, variable loadings on Factor 2, but these are still higher than on Factor 1. Only those groups with higher than intermediate loadings will be considered further. These results could not be an artifact exerted by the variations in assemblage preservation, because the CA analysis was largely based on fossil groups at generic levels, at which most specimens in all samples were identified.

Discussion

Environmental Impacts on Coccolith Formation Coccolithophores secret coccoliths inside their cells within a golgi-originated vesicle, and extrude these minute calcite plates to the outside of the cell to form exoskeleton (Westbroek et al., 1989; Pienaar, 1994). Although coccoliths retain a remarkably constant morphology within any given species (Paasche, 1968a), laboratory and field experiments on extant nannoplankton species have shown that the formation and morphology of coccoliths may be influenced to some extent by physical and chemical environmental factors. Bé and McIntyre (1965) observed that Emiliania huxleyi secretes different coccoliths in warm and cold waters. Although there exist genetic variations among different morphotypes (cryptic species) of this species (Young and Westbroek, 1991; Medlin et al., 1996), their distribution has been shown to be controlled by hydrographic conditions (e.g., Hagino et al., 2005). Laboratory experiments have shown that this species tends to produce more abnormal coccoliths, or even ceases to produce coccoliths near the extremes of its temperature growth range (Watabe and Wilbur, 1966). However, Paasche (1968b) found that the calcite produced per unit cell volume was independent of temperature. Field data from the Arafura Sea and Gulf of Carpentaria shows no relationship between the development of malformation and seawater temperate, salinity, ammonium, phosphate or dissolved oxygen level, but reveals a negative correlation with nitrate and nitrite concentrations (Okada and Honjo, 1975). This observation does not exclude salinity as a possible factor affecting calcification, as coccolithophores in the

96 southern Red Sea did not produce coccoliths during January and February when the oligotrophic surface waters have a relatively low salinity, whereas coccolith production resumed in August when nutrients and salinity reach normal conditions (Zijlstra et al., 1990). Kleijne (1990) also inferred a development of malformed coccoliths under lower salinity in the Indonesian Seas, although she cannot separate the role that nutrients play. Coccolith formation is also light- dependent (Paasche, 1964), but this process is not closely coupled with the most obvious light- requiring photosynthesis (Paasche, 1964; Herfort et al., 2002, 2004). Nutrients, including macronutrients (N, P) and micronutrients (trace elements and vitamins), have been shown to play an essential role in coccolith formation (Okada and Honjo, 1975; Paasche, 1998; Riegman et al., 2000; Shiraiwa, 2003; Rost and Riebesell, 2004), although a consensus has not been reached on the specific role of each nutrient species. Paasche (1998) reported a reduction in calcification under nitrate limitation, but Riegman et al. (2000), Shiraiwa (2003), and Rost and Riebesell (2004) observed that a limitation in nitrate and phosphate can stimulate calcification. This dichotomy seems to result from the species-specific dependence of response, as well as the complexity of response of coccolithophores to different nutrient regimes (e.g. unbalanced N/P ratios can induce coccolith deformation [Båtvik et al., 1997]). Trace elements, including magnesium, iron, copper, cadmium, and zinc, have also been documented to influence the production and calcification of coccoliths (Brand et al., 1983, 1986; Schulz et al., 2004). Mg has been shown to be essential for calcification and an abnormal concentration of this ion can cause coccolith malformation and undercalcification in Cricosphaera carterae and Emiliania huxleyi (Stillwell and Corum, 1982; Herfort et al., 2004). In addition, a combination of vitamin B12 and trace metals has been reported to induce a phase alternation between holococcolith and heterococcolith in Calyptrosphaera sphaeroidea (Nöel et al., 2004).

Knowledge of the effects of CO2 on coccolithophore calcification has also undergone recent

advances. Should an injection of CO2 into the ocean be sufficient to reduce seawater pH, the consequent acidic condition will directly impact the nature and speciation of nutrients (Turner et al., 1981; Zeebe and Wolf-Gladrow, 2001). For instance, a reduction in pH of 0.3 units could

reduce the fraction of NH3 by around 50% (Raven, 1986), while free metal elements such as iron and copper may increase, adding toxicity to the seawaters (Turner et al., 1981). It has been shown that living coccolithophores, such as Emiliania huxleyi and Gephyrocapsa oceanica, show malformation and under-calcification in their coccoliths under decreasing alkalinity

97 conditions caused by rising atmospheric pCO2 in laboratory and natural cultures (Riebesell et al., 1993, 2000; Zondervan et al., 2001; Rost and Riebesell, 2004), although its rate of photosynthesis may be elevated (Engel et al. 2005). Other calcareous organisms, such as

planktonic foraminifers and corals, also respond to increasing atmospheric pCO2 conditions by lowering their calcification rates and degrees (Gattuso et al., 1998; Bijima et al., 1999; Kleypas et al., 1999). The literature review above clearly reveals that stressed conditions, to date in terms of temperature, nutrient, and sea-water chemistry, induce deformed coccoliths, although there is no consensus on the role that a single environmental factor plays in inducing coccolith malformation. In fact, even the origin of malformation is still under debate. Abnormal coccoliths can be a product of dissolution (Roth and Berger, 1975; Schneidermann, 1977; Young and Westbroek, 1991; Young, 1994a) or teratogenesis (Okada and Honjo, 1975; Kleijne, 1990; Båtvik et al., 1997). The latter is variously referred to as “collapsed coccoliths” (Young, 1994) or “malformation” (Kleijne, 1990). These two morphological forms are readily distinguishable under the SEM (Kleijne, 1990; Båtvik et al., 1997), but such a distinction for our materials would require extensive work beyond the scope of this study, which utilizes primarily the light microscope. Our study only considers the three environmental parameters and their impacts on malformed nannofossils.

Correspondence Analysis Factors and Their Paleoenvironmental Indications Our understanding of physiological preferences of extinct nannoplankton largely comes from ecological and biogeographic studies on modern coccolithophores, as well as correlation with other paleoceanographic proxies (e.g., other microfossil groups, geochemical tracers, and sedimentologic indicators etc.), under the assumption that ecological preferences of nannoplankton taxa have not changed much over time. Research has demonstrated that coccolithophores are nearly exclusively planktonic and as a whole have a wide tolerance of temperature and salinity, although most coccolithophores thrive in warm, stratified, nutrient-poor offshore waters (Brand, 1994). This modern distribution pattern is in contrast with the cosmopolitan and diverse nature during the Cretaceous (Tappan, 1980), and was interpreted as a result of the K/T impact event which affected coccolithophores significantly, but less so diatoms (Brand, 1994; Falkowski et al., 2004). Hence, diatoms assumed dominance in nutrient-rich

98 waters previously occupied by coccolithophores, leaving the later-recovered, biologically less competitive coccolithophores relative to the newly-evolved diatoms with nutrient storage strategy, primarily to oligotrophic habitats (Brand, 1994; Falkowski et al., 2004; Katz et al., 2004). However, the temperature difference for optimum growth varies among taxa, allowing the recognition of discrete latitudinal climatic assemblages in the Atlantic and Pacific Oceans (McIntyre and Bé, 1967; McIntyre et al., 1970; Okada and Honjo, 1973). As photosynthetic organisms, coccolithophores live in the photic zone of the oceans, and the effects of light intensity lead to vertical stratification of assemblages (Okada and Honjo, 1973), and their growth and calcification sensitively respond to nutrient concentrations (Paasche, 1968a; Brand, 1994; Paasche, 1998). Because the broadly latitudinal nannoplankton temperature belts of McIntyre and Bé (1967) and Okada and Honjo (1973) coincide with oceanic divergences and central gyres, it is practically impossible to separate the effects of temperature and fertility (Bralower, 2002). A practical solution is to distinguish nannoplankton in the context of a trophic resource continuum (Bralower, 2002), based on r-K differentiation in evolutionary ecology (Young, 1994b). K-selected specialists dominate high-diversity communities in oligotrophic waters, whereas r-selected specialists constitute opportunistic species of low-diversity assemblages in eutrophic areas, with mesotrophic species between these two groups. Studies of past nannofossil biogeography have established a strong basis for physiological preference of extinct nannoplankton taxa (e.g. Haq and Lohmann, 1976; Roth and Krumbach, 1986; Watkins, 1986; Wei and Wise, 1990; Aubry, 1992, 1998; Bralower, 2002). Discoasters have been known as warm-water specialists (Bukry, 1973), but should be used with caution since there is no straightforward relationship between their abundance and water temperature (Wei and Wise, 1990; Chepstow-Lusty et al., 1992; Chapman and Chepstow-Lusty, 1997). The somewhat contradictory distribution pattern of discoasters (lower abundance at tropics than at mid latitudes as pointed out by Wei and Wise [1990]) was interpreted as the distribution of a deep dweller with an adaptation to oligotrophic conditions (Aubry, 1992) or as a suppression by higher nutrient availability (Gibbs et al., 2004). Fasciculithus and Sphenolithus are also considered to have been adapted to warm, oligotrophic waters, a finding based on their close association with Discoaster (Haq and Lohmann, 1976; Wei and Wise, 1990) and the inverse correlation with .Prinsius ٛ artini (Haq and Lohmann, 1976), a species with bipolar distribution

99 R-mode specialists indicating cool, eutrophic conditions were encountered in this study. Chiasmolithus is a well-established cool- to cold-water taxon (Wei and Wise, 1990; Firth and Wise, 1992), and Biscutum has been widely used as a productivity proxy in Cretaceous and Paleogene studies (Roth and Krumbach, 1986; Pospichal and Huber, 1992; Street and Bown, 2000; Bralower, 2002). Species of Hornibrookina have been documented most commonly at southern high latitudes (Wei and Pospichal, 1991), and their abundance pattern has been used to delineate cool-water mass boundaries across the Kerguelen Plateau in the Paleocene (Arney and Wise, 2003). At present, Coccolithus pelagicus shows an optimum growth temperature between 2 and 12ºC (Okada and McIntyre, 1979), and is well represented in North Atlantic subpolar and polar water masses (McIntyre and Bé, 1967), particularly in the Arctic-influenced Greenland Sea (Andruleit, 1997). Therefore, this taxon is widely used as a temperature and productivity proxy in Neogene and Quaternary paleoceanographic studies (McIntyre and Bé, 1967; McIntyre et al., 1970; Haq, 1980; Roth, 1994). Haq and Lohmann (1976) indicated an ecological evolution for C. pelagicus from a warmer to cooler preference during Paleogene, based on its dominance migration from the equator in the Paleocene, to mostly high latitudes in the Eocene, and then to above mid-latitudes in the Oligocene, which would otherwise suggest a gradual warming during the Paleogene. However, as all the localities dominated by this species were located in or close to divergence zones or continental shelves characterized by high nutrient availability, its migration may be alternatively considered as a response to changes in nutrient regimes. An example for this is that the high productivity under early Paleocene warm climate was signaled by a C. pelagicus bloom and confirmed by increased δ13C values (Tantawy, 2003). Hence, on this evidence C. pelagicus could be considered as an r-selected species, although this is still somewhat speculative. Toweius appears to have been cosmopolitan (Haq and Lohmann, 1976), but shows higher abundance in high latitudes and has been interpreted as eurythermal and mesotrophic (Bralower, 2002). Using the CA analysis, Factor 2 shows a high positive loading for K-mode specialist (Fasciculithus), negative loadings for r-mode specialists (Chiasmolithus, Biscutum, Hornibrookina, and Coccolithus pelagicus), and an intermediate loading for mesotrophic Toweius, suggesting an effect of surface-water fertility. Factor 1 shows a high positive loading for excursion discoasters and Markalius cf. M. apertus, an intermediate positive loading for Rhomboaster, but a high negative loading for

100 Fasciculithus and Toweius (Table 4-3, Fig. 4-5B). Excursion discoasters, including D. araneus and D. anartios, are here considered to be morphologically malformed compared to other discoasters, and which, together with Rhomboaster, have been found to form a unique, exclusive association during the PETM in the world’s oceans (Aubry, 2001; Bralower, 2002; Raffi et al., 2005), and represent a significant event in ocean surface-water history (Kahn and Aubry, 2004). As discussed above, these peculiar nannofossils mostly formed under a stressed environment. Markalius cf. M. apertus is a round or elliptical species similar to Markalius apertus but with a much larger central opening and narrower rims (Plate 4-1), and may represent a taxonomic variant subject to unusual conditions. Therefore, Factor 1 is tentatively interpreted as environmental stress associated only with surface-water chemistry and temperature, because the nature of statistical independence of the factors excludes a combination with nutrient conditions already depicted by Factor 2. Taking into account that the PETM has been attributed a massive

input of CO2 into ocean and atmosphere, Factor 2 may well represent seawater pH variations that

respond to changing seawater pCO2.

PETM Surface-water Environmental Variations Correspondence analysis suggests that 60.3% of the total information from the nannofossil assemblages represents the influence of surface-water chemistry on the evolution of ephemeral

species, suggesting that nannoplankton are sensitive to CO2-related changes in carbonate chemistry. This is because inorganic carbon assimilation by phytoplankton is largely bound to

CO2 fixation by the carboxylating enzyme, which has a low affinity to CO2 with half-saturation constants (Badger et al., 1998). Nannoplankton assemblages from Site 1259 have recorded a sharp acidification in surface waters at the onset of the PETM, followed by a step-wise gradual recovery to values lower than pre-event levels (Fig. 4-5C). This abrupt drop in seawater pH values has been attributed to the rapid release of 1500-2200 gigatons of methane (Dickens et al, 1995, 1997; Dickens, 2000). This enormous amount of methane was immediately oxidized to carbon dioxide in seawater and atmosphere, which contributed to increased temperature, reduced pH, and shoaling of the CCD and lysocline within several thousand years (Dickens, 2000; Zachos et al., 2005). An abrupt increase in surface-water fertility and consequent productivity indicated by an increasing abundance in Coccolithus pelagicus and Hornibrookina arca, preceded the dramatic

101 pH changes (Fig. 4-5C); the cause, however, remains unknown. This result agrees with the findings of Stoll and Bains (2003) using coccolith Sr/Ca ratios for Site 690 located in Weddell Sea, but contrasts with those of Bralower (2002) using nannofossil assemblage data for the same site. Theoretically, an increase of CO2 may reduce the need for CO2 concentration mechanisms

(a strategy used by many phytoplankton species to enhance supply of CO2 for photosynthesis), and thus lower the metabolic costs to the cells for inorganic carbon sequestration (Engel et al., 2005). This is hypothesized to increase the production of extracellular organic matter (Engel,

2002). In practice, an elevated CO2 concentration can lead to a shift in phytoplankton composition due to competition (Tortell et al., 2002), as with the Coccolithus pelagicus/Toweius abundance switch observed in this study. Variations in productivity would have the potential to change phytoplankton composition. Furthermore, Site 1259 is located close to South American drainage systems, where today water flowing into the Atlantic is deflected north and delivered to this site. During the PETM, the assumed humid climate and high runoff (Bowen et al., 2004) would deliver large amounts of nutrients and create a fertile condition at this site. Lastly, the PETM induced a much higher SST increase at high latitudes than in the tropics (Kennett and Scott, 1991; Zachos et al., 1993; Zachos et al., 2003), which would have affected the former region more significantly given the presumably higher annual average SST in the latter region. Therefore, oligotrophy in open-ocean circum-Antarctic waters as suggested by Bralower (2002) does not necessarily imply worldwide oligotrophy, particularly in continental margin settings as recently pointed out by Gibbs et al. (2006). These authors contrasted the mesotrophic conditions that existed across the PETM off New Jersey (USA) with those of the west-central paleoequatorial Pacific. Our observations support and reinforce those by Gibbs et al. (2006).

Response of Nannoplankton to the PETM “Excursion taxa” is the term applied by Kelly et al. (1996) to the ephemeral foraminiferal ecophenotypes restricted to the interval of the CIE, including species of the genera Acarinina and Morozovella. This term was also used to describe the short-lived discoasters D. anartios and D. araneus that evolved at the P/E boundary (e.g., Bybell and Self-Trail, 1995; Kelly et al., 1996; Agnini et al., 2004). Rhomboaster was later added due to its close association with D. araneus (Kahn and Aubry, 2004). The most prominent feature of the discoaster excursion taxa is the irregular arrangement and length of their rays, which is considered here as a malformation

102 compared with other normally-developed, symmetrical discoasters. In our study, we also find a high abundance of Markalius cf. M. apertus associated with these other, better-known excursion taxa. It is generally accepted that the PETM extreme environmental changes caused the nannoplankton response (producing malformed coccoliths) due to their concomitant occurrence. What individual physical factor accounts for the response is the subject of current debate. Temperature is not likely to be the cause since excursion taxa are present in sites where there is only limited temperature increase (Kelly et al., 1996). Oligotrophy is a possible alternative, and recent studies have shown the effects of nutrients (N, P) not only on phytoplankton cell number and size, but also coccolith number and calcification rate (Båtvik et al., 1997; Paasche, 1998; Riegman, 2000; Shiraiwa, 2003), as reviewed above. However, this suggestion cannot explain the widely-distributed occurrence of excursion taxa where nutrient conditions are variable. By process of elimination, therefore, we believe a decrease in pH values is one of the most likely factors for causing malformation of the coccoliths. Seawater pH values have been shown, in cultures as well as in field, to reduce calcite production and to increase malformed coccoliths in modern dominant coccolithophores Emiliania huxleyi and Gephyrocapsa oceanica (Riebesell et al., 2000; Zondervan et al., 2001, 2002; Engel et al., 2005). In the culture studies and field experiments by Riebesell et al. (2000), Emiliania huxleyi grew massive coccoliths, while Gephyrocapsa oceanica developed coccoliths with an incomplete central structure under elevated CO2 concentrations, which could explain the cases such as in Rhomboaster and Markalius cf. M. apertus. Rhomboaster is considered to be massive form relative to discoasters, given its larger mean mass of calcite relative to discoasters as estimated here using the methods of Young and Ziveri (2000) and by using shape factors for morphologically similar extant species. The reduction of abundance in Toweius and Fasciculithus, by ~12% and ~35%, respectively, was mainly picked up by Markalius cf. M. apertus and excursion taxa with an increase by ~25% and ~12%, respectively. The acme of Markalius cf. M. apertus at the PETM has not been reported elsewhere, although Gibbs et al. (2006) mentioned a minor peak of Markalius (species unspecified) at their open ocean Site 1209. Based on our limited data, we speculate that Markalius cf. M. apertus is a local opportunistic species that has taken advantage of the extreme conditions and responded sensitively to the PETM. It is apparently not as widely distributed as

103 the previously defined excursion taxa, however, and persists higher in the section than the others although in lower abundance (Fig. 4-2). It appears to be corroded rather than malformed. Its high abundance at the PETM at our Site 1259 may simply be in response to an influx of nutrients and/or lower salinity induced by increased riverine runoff at this continental-margin locality.

Conclusions

Quantitative investigation of samples from ODP Site 1259 has documented fundamental changes in the calcareous nannofossil assemblages across the PETM. Dominance of relative abundance switched from K-mode specialists, such as Discoaster, Fasciculithus and Toweius, to excursion taxa and r-mode specialists, such as Coccolithus pelagicus, Hornibrookina arca, and Chiasmolithus. Interpretation of the two extracted factors from correspondence analysis based on the ecology and morphology of nannofossil taxa reveals an abrupt drop in seawater alkalinity preceded by an increase in surface-water productivity. Decreasing seawater pH might have caused the emergence of excursion taxa. Contrasting results between this study with that of Bralower (2002) in the Southern Ocean can probably be attributed to the differences of locality and subsequent differences in oceanographic conditions, as demonstrated by Gibbs et al. (2006) between the continental margin setting of eastern North America and the central paleoequatorial Pacific Ocean. Our study shows that mesotrophic conditions existed along the equatorial Atlantic continental margin of South America as well. Further work needed to done to test the explanations provided here is to calculate coccolith accumulation rates using absolute abundances instead of relative percentage abundances. This would allow better estimate variations in nannoplankton calcification and provide further evidence for surface-water changes. Studying the size variations in specimens of Coccolithus pelagicus could also reveal additional information.

Acknowledgements

We thank all the participants of ODP Leg 207 for their contributions relevant to this paper. Special thanks go to Dr. Dick Norris for providing samples from the U-channel through the

104 archive half of the core; he also kindly provided use of his unpublished bulk carbonate isotope data on the materials. We thank two anonymous reviewers for constructive and helpful comments that greatly improved this manuscript. In addition, members of the Calcareous Nannofossil Laboratory at Florida State University, especially Denise Kulhanek, provided helpful comments and assistance. Dr. William Parker (FSU) kindly provided expertise in adapting the statistical program, and Cinthia Wise assisted with the French translation. This research used samples and data provided by ODP. ODP is sponsored by the U.S. National Science Foundation (NSF) and participating countries under management of Joint Oceanographic Institutions (JOI), Inc. Direct funding for this research was provided by a JOI USSAC grant (Task Order T308A33) to SWW. General laboratory support was provided by NSF-OPP grants 0126218 and 0230469.

105

Figure 4-1. Map showing the location of ODP Site 1259 superimposed on a 55.5 Ma plate reconstruction.

106

Figure 4-2. Percentage abundance of selected calcareous nannofossil taxa and bulk carbonate δ13C (R.D. Norris, unpublished data, 2006) across the PETM from ODP Site 1259.

107

Figure 4-3. Close-up photograph of the P/E boundary recovered in ODP Hole 1259B (from Shipboard Scientific Party, 2004b).

108

Figure 4-4. Variations of nannofossil assemblage preservation and diversity across the PETM from ODP Site 1259.

109

Figure 4-5. Results of correspondence analysis. A. Percent variance of each factor; B. Factors 1 and 2 cross plot; C. Sample scores on Factors 1 and 2, indicating seawater pH and fertility variations.

110 Table 4-1. Calcareous nannofossil distribution chart and counts in ODP Hole 1259B.

111 Table 4-1 – continuded

112 Table 4-1 – continuded

113 Table 4-2. List of calcareous nannofossil taxa identified.

Group Taxa Group Taxa Biscutum Biscutum sp. Toweius Toweius eminens Chiasmolithus Chiasmolithus bidens Toweius occultatus Chiasmolithus consuetus Toweius pertusus Chiasmolithus solitus Toweius sp. Chiasmolithus sp. Other taxa Braarudosphaera bigelowii Coccolithus palagicus Coccolithus pelagicus Campylosphaera dela Excursion taxa Discoaster anartios Campylosphaera eodela Discoaster araneus Coccolithus formosus Other discoasters Discoaster cf. falcatus Coccolithus subpertusus Discoaster cf. mohleri Coronocyclus prionion Discoaster diastypus Cruciplacolithus cribellum Discoaster falcatus Cruciplacolithus latipons Discoaster lenticularis Ellipsolithus distichus Discoaster mohleri Ellipsolithus macellus Discoaster multiradiatus Lophodolithus nascens Discoaster salisburgensis Markalius apertus Discoaster sp. Markalius inversus Fasciculithus Fasciculithus involutus Neochiastozygus concinnus Fasciculithus sp. Neochiastozygus distentus Fasciculithus thomasii Neochiastozygus junctus Fasciculithus tympaniformis Neococcolithes dubius Hornibrookina Hornibrookina arca Neococcolithes protenus Markalius cf.M. apertus Markalius apertus Placozygus sigmoides Rhomboaster Rhomboaster cuspis Pontosphaera sp. Rhomboaster spineus Rhabdosphaera tenuis Rhomboaster sp. Scapholithus fossilis Tribrachiatus bramlettei Scapholithus rhombiformis Sphenolithus Sphenolithus anarrhopus Transversopontis sp. Sphenolithus moriformis Thoracosphaera sp. Zygodiscus adamas Zygrablithus bijugatus

114 Table 4-3. Results of correspondence analysis of relative abundance data.

Group Factor 1 Factor 2 Factor 3 Factor 4 Factor 5 Biscutum 0.008 -0.060 -0.101 0.152 -0.322 Chiasmolithus -0.037 -0.058 -0.096 -0.003 0.093 Coccolithus pelagicus -0.003 -0.295 0.096 0.044 -0.094 Excursion taxa 0.523 0.347 0.146 -0.510 -0.366 Other Discoaster -0.032 -0.159 0.436 -0.042 -0.124 Fasciculithus -0.324 0.792 -0.191 0.104 -0.054 Hornibrookina arca -0.032 -0.245 -0.748 -0.071 -0.458 Markalius cf. apertus 0.574 0.122 -0.056 0.131 0.053 Rhomboaster 0.283 0.157 0.045 0.717 -0.090 Sphenolithus 0.114 -0.047 -0.090 -0.272 0.295 Toweius -0.397 0.134 0.139 -0.131 0.005 Zygodiscids 0.179 0.059 -0.362 -0.117 0.620 Other taxa 0.073 -0.099 -0.014 0.245 0.188 Percent variance 60.294 24.294 3.6661 2.5398 2.0147 Cumulative percent 60.294 84.588 88.2541 90.7939 92.8086

115

Plate 4-1. Calcareous nannofossils across the PETM in ODP Hole 1259B.

116 Scale bar equals 3 microns (µm). SEM = scanning electronic microscope, XP = cross-polarized light, TL = transmitted light. All specimens are from Sample 1259B-8R-4, 51-52 cm (367.71 mbsf) except otherwise stated.

1-5, 11-21. Discoaster araneus 1-5. SEM, Zone NP10a 11-21. TL, Zone NP10a 6-9, 22-26. Markalius cf. Markalius apertus 6-9. SEM, Zone NP10a 22. TL, Zone NP10a 23. Same specimen, XP 24-26. XP, Zone NP10a 10. Tribrachiatus bramlettei SEM, Zone NP10a 27. Rhomboaster sp. XP, Zone NP10a 28. Rhomboaster spineus TL, Zone NP10a, Sample 1259B-8R-4, 83-84 cm (368.03 mbsf) 29. Toweius sp. SEM, Zone NP10a 30. Toweius cf. T. selandianus SEM, Zone NP10a 31-32. Toweius eminens 31. XP, Zone NP10a 32. XP, Zone NP10b, Sample 1259B-8R-2, 31-32 cm (364.51 mbsf) 33. Toweius cf. eminens XP, Zone NP10b, Sample 1259B-8R-2, 31-32 cm (364.51 mbsf) 34. Toweius pertusus XP, Zone NP10a 35-36. Coccolithus pelagicus XP, Zone NP10b, Sample 1259B-8R-3, 20-21 cm (365.90 mbsf) 37-38. Markalius inversus 37. XP, Zone NP10b, Sample 1259B-8R-3, 20-21 cm (365.90 mbsf) 38. XP, Zone NP10a 39-40. Markalius apertus 39. XP, Zone NP10b, Sample 1259B-8R-2, 31-32 cm (364.51 mbsf) 40. XP, Zone NP10b, Sample 1259B-8R-3, 20-21 cm (365.90 mbsf)

117

Plate 4-2. Calcareous nannofossils across the PETM in ODP Hole 1259B.

118 Scale bar equals 3 microns (µm). SEM = scanning electronic microscope, XP = cross-polarized light, TL = transmitted light.

1-2. Chiasmolithus bidens 1. XP, Zone NP9a, Sample 1259B-8R-4, 148-149 cm (368.68 mbsf) 2. XP, Zone NP10b, Sample 1259B-8R-3, 20-21 cm (365.90 mbsf) 3-5. Chiasmolithus consuetus 3, 5. XP, Zone NP9a, Sample 1259B-8R-4, 148-149 cm (368.68 mbsf) 4. XP, Zone NP10b, Sample 1259B-8R-2, 31-32 cm (364.51 mbsf) 6. Chiasmolithus solitus XP, Zone NP10b, Sample 1259B-8R-2, 31-32 cm (364.51 mbsf) 7-9. Chiasmolithus californicus 7. XP, Zone NP9a, Sample 1259B-8R-4, 148-149 cm (368.68 mbsf) 8-9. XP, Zone NP10b, Sample 1259B-8R-3, 20-21 cm (365.90 mbsf) 10. Cruciplacolithus sp. XP, Zone NP10b, Sample 1259B-8R-3, 20-21 cm (365.90 mbsf) 11. Cruciplacolithus latipons XP, Zone NP10b, Sample 1259B-8R-2, 31-32 cm (364.51 mbsf) 13-15. Discoaster multiradiatus 13. SEM, Zone NP10a, Sample 1259B-8R-4, 51-52 cm (367.71 mbsf) 14. TL, Zone NP10a, Sample 1259B-8R-4, 51-52 cm (367.71 mbsf) 15. TL, Zone NP10b, Sample 1259B-8R-3, 20-21 cm (365.90 mbsf) 16. Discoaster lenticularis TL, Zone NP10a, Sample 1259B-8R-4, 51-52 cm (367.71 mbsf) 17. Discoaster sp. TL, Zone NP9a, Sample 1259B-8R-4, 148-149 cm (368.68 mbsf) 18. Fasciculithus tympaniformis XP, Zone NP9a, Sample 1259B-8R-4, 148-149 cm (368.68 mbsf) 19. Fasciculithus thomasii XP, Zone NP9a, Sample 1259B-8R-4, 148-149 cm (368.68 mbsf) 20. Fasciculithus sidereus XP, Zone NP10b, Sample 1259B-8R-3, 20-21 cm (365.90 mbsf) 21-25. Campylosphaera dela 21. XP, Zone NP10b, Sample 1259B-8R-2, 1-2 cm (364.21 mbsf) 22-25. XP, Zone NP10b, Sample 1259B-8R-3, 20-21 cm (365.90 mbsf) 26-27. Campylosphaera eodela XP, Zone NP10b, Sample 1259B-8R-2, 31-32 cm (364.51 mbsf) 28. Neochiastozygus junctus XP, Zone NP10b, Sample 1259B-8R-3, 20-21 cm (365.90 mbsf) 29. Neococcolithes protenus XP, Zone NP9a, Sample 1259B-8R-4, 148-149 cm (368.68 mbsf) 30. Placozygus sigmoides XP, Zone NP10a, Sample 1259B-8R-4, 51-52 cm (367.71 mbsf) 31-32. Hornibrookina arca 31. XP, Zone NP10b, Sample 1259B-8R-3, 20-21 cm (365.90 mbsf) 32. XP, Zone NP10b, Sample 1259B-8R-2, 31-32 cm (364.51 mbsf)

119 33-36. Biscutum sp. XP, Zone NP10b, Sample 1259B-8R-3, 20-21 cm (365.90 mbsf) 37. Ellipsolithus macellus XP, Zone NP10b, Sample 1259B-8R-2, 31-32 cm (364.51 mbsf) 38. Ellipsollithus distichus XP, Zone NP10b, Sample 1259B-8R-3, 20-21 cm (365.90 mbsf) 39. Thoracosphaera sp. XP, Zone NP10a, Sample 1259B-8R-4, 51-52 cm (367.71 mbsf) 40. Coronocyclus prionion XP, Zone NP10b, Sample 1259B-8R-2, 31-32 cm (364.51 mbsf) 41. Scapholithus rhombiformis XP, Zone NP9a, Sample 1259B-8R-4, 148-149 cm (368.68 mbsf) 42. Pontosphaera scissura XP, Zone NP9a, Sample 1259B-8R-4, 148-149 cm (368.68 mbsf)

120

CHAPTER 5

TAXONOMIC NOTE: A NEW COCCOLITHUS SPECIES THRIVING DURING THE PALEOCENE/EOCENE THERMAL MAXIMUM

Introduction

A high-resolution calcareous nannofossil biostratigraphic study of clay-rich materials from Hole 1259B, Ocean Drilling Program (ODP) Leg 207 to the Demerara Rise, southwest equatorial Atlantic (Shipboard Scientific Party, 2004b), yielded a new species, Coccolithus bownii. Based on a similar variation in abundance pattern compared to that of “excursion taxa” (e.g., Discoaster araneus), this is considered to be an opportunistic species (Jiang and Wise, 2006) that took advantage of the extreme environmental changes associated with the Paleocene/Eocene Thermal Maximum (PETM) and thrived in stressed environments. Similar acmes have also been observed in taxa attributed to Coccolithus subpertusus in several Egyptian PETM sections (e.g., Dupuis et al., 2003; Knox et al., 2003). These may be of C. bownii instead (taking into account the close resemblance between these two species), which would render this new species an important, widespread ecological indicator of stressed environments.

Material and Methods

The materials examined were sampled by a U-channel through the archive halves of ODP Sections 1259B-8R-2 and 8R-4. Sections 8R-2 and -3 consist primarily of light greenish nannofossil chalk with foraminifers, whereas Section 8R-4 contains a dark yellowish brown claystone interbedded within the nannofossil chalk (Shipboard Scientific Party, 2004b).

121 Preparation of smear slides followed standard techniques (Bown and Young, 1998), using Norland-61 optical adhesive as a permanent mounting medium. Slides were analyzed using a Zeiss Axioskop II microscope under cross-polarized light (XPL), transmitted light (TL), and phase-contrast light (PC) at 1000x magnification. An FEI Nova 400 Nano scanning electron microscope (SEM) was employed to examine fine-scale structures. A settling technique was used in stub preparation to remove large particles and fine clays, whereas a freshly-broken fracture section was also coated to provide pristine, clean clay-free specimens. All materials are archived in the collections of the Calcareous Nannofossil Laboratory at the Department of Geological Sciences, Florida State University (FSU), Tallahassee, Florida, USA.

Systematic Paleontology

Family COCCOLITHACEAE Kamptner, 1928 Genus Coccolithus Schwarz, 1894 Coccolithus bownii, Jiang and Wise, sp. nov. (Pl. 5.1, Figs. 1-21)

Markalius sp. cf. M. apertus, Jiang and Wise, 2006, pp. 236-237, pl.1, figs. 6-9, 22-26. Diagnosis: A nearly round Coccolithus with a large central opening. Description: The somewhat elliptical placolith consists of two widely separated shields (Figs. 3-5), each composed of ~44-48 imbricate elements (Figs. 3, 4, 14, and 15). In distal view, the distal shield is dextrogyre with its inner margin covered by a shingling of thin central-area laths (Figs. 14-16). The proximal shield consists of two closely appressed and superimposed cycles of elements that both extend out to its margin, the proximal-most being strongly dextrogyre in proximal view (Figs. 1, 4, and 5). This structure of upper and lower layers of the proximal shield, though somewhat difficult to discern in our specimens, is strongly reminiscent of that shown in artificially etched specimens of Coccolithus pelagicus (Wallich) Schiller (1930) by Young (1992, p. 4, figs. 1-5). The central opening is large, occupying about 40-50% of the coccoliths diameter along the minor axis (Figs. 1-21). Only the proximal shield and laths appear bright in polarized light

122 (Figs. 17-19), producing a herringbone birefringence pattern characteristic of the genus. In etched specimens, the birefringence pattern does not change except that the central opening appears much larger. Differentiation: Coccolithus bownii resembles in outline C. occidentalis Black, n. comb. However, the latter species has a striking chevron-pattern of elements formed by the two cycles of elements of the proximal shield, of which the proximal-most is only half the width of the underlying cycle. The central-area opening is also considerably smaller, occupying only a quarter of the total width measured along the minor axis (see Black, 1964, pl. 52, fig. 1). The large central opening and somewhat elliptical outline differentiate C. bownii from other Coccolithus species, most of which that are more elliptical. Coccolithus formosus, on the other hand is nearly circular, but has a much smaller central opening (about 20% or less of the total minor-axis diameter; see Black, 1964, pl. 50, fig. 2 [as “C. lusitanicus”]). Remarks: As pointed out above, construction of the proximal shield of C. bownii resembles closely that of the modern C. pelagicus, i.e., it consists of two thin, closely appressed, superimposed layers of near equal width. The presence of this construction in the lower Eocene shows that this feature occurs in the older as well as the younger parts of the Coccolithus lineage. This construction contrasts with other contemporaneous forms such as C. occidentalis Black, n. comb., and C. muiri (= “E. ovalis”; see Roth, 1970, p. 841), which Black (1964) placed in a separate genus, Ericsonia, based on the much shorter widths of the proximal-most layer of the proximal shield. The relative widths of those two layers, however, vary significantly among species and over time (Wise, 1983, p. 505), hence it is difficult to draw a non-arbitrary dividing line between the genus Coccolithus and Black’s Ericsonia. Following the lead of Bukry (1973) and Wise (1973, 1983), we continue to place all such species in the genus Coccolithus. Deviation of name: After Dr. Paul Bown (University College London) in honor of his many valuable contributions to this subdiscipline of micropaleontology. Occurrence: Coccolithus bownii is common to abundant within Zone NP10 of Martini (1971) in ODP Hole 1256B. Specifically, it shows an acme within the PETM interval, and decreases in abundance thereafter upsection (Jiang and Wise, 2006, fig. 2). Should the acmes of C. subpertusus in Egypt actually be of C. bownii, then this new species would have at least a Tethyan distribution. Dimensions: 6-8 μm (holotype: 6.2 μm).

123 Holotype: Pl. 5.1, Fig. 1 Paratype: Pl. 5.1, Figs. 3, 4, 8, 10, 14, and 15. Type locality: ODP Hole 1259B on Demerara Rise, South America (9°18.0485´N, 54°11.9448´W), Sample 1259B-8R-4, 57-58 cm.

Coccolithus insolitus (Perch-Nielsen, 1971) Jiang, Ladner and Wise, n. comb.

Basionym: Ericsonia insolita, Perch-Nielsen, 1971, Elektronenmikroskopische Untersuchungen an Coccolithen und verwandten Formen aus dem Eozän von Dänemark. Det Kongelige Danske Videnskabernes Selskab Biologiske Skrifter, v. 18, p. 13, pl. 1, fig. 1; pl. 7, figs. 4, 6; pl. 61, figs. 14, 15. Remarks: See remarks under C. bownii. This combination was previously made by Ladner and Wise in Wise et al. (2002) in a publication distributed via CD-ROM only. New combinations published in electronic media are invalid under provisions of the International Code of Botanical Nomenclature published in 2000 (Article 29.1; Greuter et al., 2000).

Coccolithus occidentalis (Black, 1964) Jiang, Ladner and Wise, new comb. Basionym: Ericsonia occidentalis, M. Black, 1964, Cretaceous and Tertiary coccoliths from Atlantic seamounts, Palaeontology, v. 7(2), pl. 52, figs. 1, 2, pp. 311-312. Remarks: See remarks under C. bownii. This combination was previously made by Ladner and Wise in Wise et al. (2002) (invalid; see remarks under C. insolitus).

Acknowledgements

We thank Kim Riddle and Thomas Fellers of the FSU Biology Unit 1 SEM laboratory for helpful assistance and instruction on the operation of the newly installed FEI Nova 400 Nano SEM, which was purchased through a grant from the U.S. National Science Foundation (NSF grant 0520950). We are grateful to ______(reviewers) for constructive and helpful comments that greatly improved this manuscript. The samples were provided by the Ocean Drilling Program (ODP). ODP is sponsored by the US National Science Foundation (NSF) and participating countries under management of Joint Oceanographic Institutions (JOI), Inc. General

124 support for the Calcareous Nannofossil Laboratory in the FSU Department of Geological Sciences was provided by NSF-OPP grants 0125526 and 0230469 to SWW.

125

Plate 5-1. Coccolithus bownii Jiang and Wise, 2007. Sample ODP 206-1256B-6H-2, 115-117 cm. Scale bar = 2 μm. SEM = scanning electronic microscope; XPL = cross-polarized light; TL = transmitted light; PC = phase contrast.

126

CHAPTER 6

ABRUPT TURNOVER IN CALCAREOUS-NANNOPLANKTON ASSEMBLAGES ACROSS THE PALEOCENE/EOCENE THERMAL MAXIMUM: IMPLICATIONS FOR SURFACE-WATER OLIGOTROPHY OVER THE KERGUELEN PLATEAU, SOUTHERN INDIAN OCEAN

Abstract

Ocean Drilling Program (ODP) Core 183-1135A-25R-4 from the Kerguelen Plateau in the Indian Ocean sector of the Southern Ocean represents only the second complete, expanded section of the Paleocene/Eocene Thermal Maximum (PETM; ~55 Ma) recovered from Antarctic waters. Calcareous nannoplankton at this site underwent an abrupt, fundamental turnover across the PETM as defined by a carbon isotope excursion. Although Chiasmolithus, Discoaster, and Fasciculithus exponentially increase in abundance at the onset, the former abruptly drops but then rapidly recovers, whereas the latter two taxa show opposite trends due to surface-water oligotrophy. These observations confirm previous results from ODP Site 690 on Maud Rise. The elevated pCO2 that accompanied the PETM caused a shoaling of the lysocline and carbonate compensation depth, leading to intensive dissolution of susceptible holococcoliths and poor preservation of the assemblages. Similarities and contrasts between the results of this study and previous work from open-ocean sites and shelf margins further demonstrate that the response to the PETM was consistent in open-ocean environments, but could be localized on continental

127 shelves where nutrient regimes depend on the local geologic setting and oceanographic conditions.

Introduction

One of the major paleoclimate discoveries of the scientific Ocean Drilling Program (ODP) was the Paleocene/Eocene Thermal Maximum (PETM) at Maud Rise (ODP Site 690) in the Southern Ocean (SO; Kennett and Stott, 1991; Bralower, 2002). We report here a second such site located half way around the Antarctic continent in the SO at ODP Site 1135 on the Kerguelen Plateau (KP; Fig. 6-1). The PETM was a catastrophic, rapid greenhouse-forced global warming event ~55 m.y. ago (surface- and bottom-water temperatures rose by ~5ºC at high latitudes within <22 k.y.; [Kennett and Stott, 1991; Zachos et al., 2003; Norris and Röhl, 1999]). The environmental perturbations associated with the PETM triggered an abrupt overturn in ocean chemistry (Dickens et al., 1995; Zachos et al., 2005) and circulation (Kennett and Stott, 1991; Numes and Norris, 2006). The concomitant biotic turnovers include a mass extinction (30%–50%) of benthic foraminifers, an evolutionary radiation of land mammals and vegetation, a diversification of the planktonic foraminifers, as well as the evolution of ephemeral “excursion taxa” among planktonic foraminifera and nannofossils (see Kelly et al. [1996], Aubry et al. [1998] and references therein). A global, negative carbon-isotope excursion (CIE) of ~3‰ that accompanied the PETM is most parsimoniously explained by a massive release of sedimentary methane hydrates to the ocean-atmosphere systems (Dickens et al., 1995). Alternatively, the carbon that fueled the PETM may have been derived the thermal cracking of methane in coals in the North Atlantic (e.g., Svensen et al., 2004; Storey et al., 2007). Either way, methane, or its subsequent oxidation product, carbon dioxide, would have immediately contributed to greenhouse warming and changes in ocean chemistry. Despite its global extent, sedimentary archives of the PETM are limited and imperfect. Low- latitude PETM records generally contain a dissolution interval barren of calcareous fossils; hence we cannot say from those nannofossil records what happened during the first a few k.y. of the event. Our high-latitude records, however, have yielded calcareous fossils, although at ODP Site

128 1135 they are altered by diagenesis (e.g., fragmentation, dissolution, and overgrowth) that to some degree hinders the correct identification of species and even genera. Surface-dwelling plankton appear to have responded to the PETM environmental changes in fundamentally different ways from benthic organisms, but their strategies remain poorly understood due to the limited number of high-resolution investigations using calcareous nannofossils (e.g., Monechi et al., 2000; Bralower, 2002; Gibbs et al., 2006) and geochemical proxies (e.g., Sr/Ca ratios and biogenic Barium accumulation rates [Stoll and Bains, 2003; Bains et al., 2000]). Contrasts among previous studies in the equatorial Atlantic, SO, New Jersey continental margin, and central paleoequatorial Pacific (Fig. 6-1) suggest that local differences in geologic setting and oceanographic conditions influence the biotic response to the PETM (Bralower, 2002; Gibbs et al., 2006; Jiang and Wise, 2006). Controversy on the phytoplankton’s response to the PETM even arises from different productivity proxies at the same locality on Maud Rise in the SO (Bains et al., 2000; Bralower, 2002; Stoll and Bains, 2003). In order to determine the actual impact of the PETM on surface-water dwellers at high southern latitudes, we evaluated high- resolution variations in nannofossils and bulk stable isotopes at Site 1135 in 1567 m of water on the KP (Fig. 6-1), a sequence that was first studied in a low-resolution study by Arney (2002).

Materials and Methods

ODP Core 183-1135A-25R-4 is a carbonate-rich nannofossil and foraminifer ooze, selected for this study because of the coherence and completeness of the portion recovered. Another PETM section was recovered farther south on the KP at Site 738, but recovery was not as complete and no detailed assemblage work has been done on that section (Bralower, 2007, pers. comm.). Samples were taken at a 2-cm spacing throughout the core. Slide preparation followed the modified “glass-bead” technique of Okada (1992), whereas taxonomic concepts follow Bown (2005); techniques to determine percentage abundance were adapted from Jiang and Wise (2006). Assemblage preservation was quantified using the ratio of identifiable specimens to indeterminate isolated nannofossil shields. Five to ten bulk foraminifers from the 63-μm fraction were analyzed on a Finnigan Mat Delta plus XP mass-spectrometer in the National High Magnetic Field Laboratory (NHMFL) at FSU.

129 Results: Variations in Nannofossil Assemblages

ODP Core 183-1135A-25R-4 represents a stratigraphically complete, expanded section across the PETM as defined by the extent of the CIE (Fig. 6-2), and contains poorly preserved nannofossil assemblages. The severe diagenetic alteration renders most specimens identifiable only at the generic level, although important ecological indicators such as Chiasmolithus, Discoaster, Fasciculithus, Sphenolithus, and Zygrhablithus, survived due to their robust construction. Biscutum and Hornibrookina are also readily distinguishable when present. Unidentifiable isolated nannofossil shields, and to a lesser extent, Toweius, Coccolithus pelagicus, and Zygrhablithus bijugatus dominate assemblages (Fig. 6-2). The isolated shields are quite abundant at the onset of the PETM, drastically decrease after the CIE peak, increase somewhat during the CIE’s recovery, then decrease upsection. Z. bijugatus shows a steady decrease in abundance from the onset to the peak of the CIE, then abruptly increases thereafter, but remains high and increases again at the top of the section. The ecological indicators demonstrate significant stratigraphic variations. Discoaster, Fasciculithus, and Sphenolithus exponentially increase in abundance from the onset of the PETM, and peak prior to the CIE peak, at which point they decrease abruptly and then remain consistently low thereafter. The abundance of Chiasmolithus also demonstrates an exponential increase at the onset, abruptly drops to its lowest value prior to the CIE peak, then mimics the exponential recovery of the CIE, but decreases again before CIE’s full recovery. Biscutum and Hornibrookina occur sporadically in the section, present with abundant Chiasmolithus but at low levels or absent when Chiasmolithus abundance is low.

Discussion

Stable isotope pattern The onset of the CIE occurs within an 18-cm interval before the peak is reached (Fig. 6-2). In most deep-sea sections, this onset usually occurs in 1 or 2 cm of core. This, we believe, is the result of high sedimentation rates at this relatively shallow oceanic site, as evidenced by the high numbers of dissolution-susceptible holococcoliths (Z. bijugatus) preserved throughout the sequence. Thus, this sequence appears to be ideal for high-resolution studies. The full course of

130 the CIE exhibits a more or less symmetrical “V” shape in our data, indicating that the release of light carbon was a gradual, single injection, instead of multiple pulses as suggested by Bains et al., 1999.

Ecological implications of assemblage variations Covariations among nannofossil abundances and their corresponding correlation coefficients highlight taxa with similar trends, suggesting likely ecological groupings and possibly similar preservation states (Fig. 6-2, Table 6-1). Based on their biogeography and relationships with paleoenvironmental proxies, Discoaster, Fasciculithus, and Sphenolithus are considered as K- selected specialists adapted to warm, oligotrophic waters (e.g., Haq and Lohmann, 1976; Wei and Wise, 1990; Chepstow-Lusty et al., 1992); whereas r-mode taxa include Chiasmolithus, Biscutum, and Hornibrookina thriving in cool, eutrophic conditions (e.g., Wei and Wise, 1990; Wei and Pospichal, 1991; Bralower, 2002). Nannofossil assemblages show significant variations within the section. These variations cannot be an artifact of selective dissolution, based on the fact that 1) there is a high negative correlation between K- and r-selected taxa in the lower section with its low abundances of Z. bijugatus (a dissolution-susceptible holococcolith), and 2) r-selected taxa fluctuate while K-selected taxa do not in the upper section where Z. bijugatus consistently increases (Fig. 6-2). Otherwise there would be a consistent relationship between these two groups. Variations in preservation may in part account for the fluctuations; however, environmental factors are likely responsible for the abrupt changes across the PETM. At Site 1135, abundances of K-selected taxa covary and show opposite tends to r-mode specialists (Fig. 6-2). Although both nutrient availability and temperature exerted an influence on the abundance increases in these K-selected taxa prior to the peak of the PETM, their abrupt decrease thereafter coincides with a warming trend of 8ºC in the high latitudes (Kennett and Stott, 1991). This suggests that nutrient availability exerted a stronger influence at this site. As Z. bijugatus, a readily distinguishable holococcolith even in poorly preserved materials, is susceptible to dissolution (see Crudeli et al., 2006 and references therein), it is therefore a sensitive indicator of seawater pH values induced by the rapid release of 1500-2200 gigatons of methane (Dickens et al., 1995). The consequent effects of corrosion diminished nannofossil preservation, raising the number of indeterminate isolated nannofossil shields and fragments

131 during the development of the CIE. These decreased, however, as the system reverted to pre- event conditions.

Nannoplankton productivity changes across the PETM in the Southern Ocean The variations in nannofossil assemblages at Site 1135 primarily reflect surface-water productivity presumably with little changes in salinity as inferred from the complementary SO Site 690 (Kennett and Stott, 1991). Despite a progressive warming towards the peak of the PETM inferred by decreasing δ18O values, the overall assemblage changes do not reflect a steady trend toward oligotrophy, but rather signify eutrophication at the onset as indicated by the exponential increase of r-mode taxa in contrast to the negligible increase of k-selected specialists (Fig. 6-2). Similar precursor events just prior to the onset of the PETM were also observed at ODP Sites 690 and 1259 (Bralower, 2002; Jiang and Wise, 2006, respectively); the cause, however, remains unknown, although this eutrophic condition could be a spontaneous response of the destabilized ocean circulation system upon reaching or crossing its threshold value. Thereafter, the abrupt decrease in abundance of r-selected taxa and concomitant increase of K-selected specialists denotes a fundamental oceanographic turnover that produced an oligotrophic condition. This result is quite consistent with that from Site 690 (Bralower, 2002), suggesting that surface-water oligotrophy prevailed in the southern high latitudes. Contrasts between these results with other studies at Site 690 using geochemical and benthic assemblages have already been addressed in other recent studies (see Gibbs et al. [2006] and references therein).

Adaptive response of nannoplankton to the PETM “Excursion taxa” (ephemeral ecophenotypes restricted to the CIE; Kelly et al., 1996; Kahn and Aubry, 2004) have been interpreted as malformed nannofossils (see Jiang and Wise [2006] and references therein). Their sudden occurrence coincident with the CIE is considered to be a result of environmental-stress forcing (e.g., Jiang and Wise, 2006). Although a global distribution for these excursion taxa has been suggested (Kahn and Aubry, 2004; Raffi et al., 2005), this appears to be true only for Discoaster araneus because the other excursion taxa have not been reported in the high southern latitudes (Sites 690 and 1135). The percent abundance of D. araneus decreases markedly from the low to high latitudes, but could not be entirely

132 temperature-controlled, taking into account rather flat equator-pole temperature gradient during the PETM and the fact that non-excursion discoasters show similar abundances at both low and high latitudes (Raffi et al., 2005; Jiang and Wise, 2006). The cause, however, still remains unknown due to limited data. At Site 1135, nannofossil abundance patterns changed in step with changes in seawater chemistry during the development of the CIE, which further confirms that K- and r-selected taxa are sensitive environmental indicators (e.g., Bralower, 2002). However, their extreme values preceded the peak of the CIE, and they recovered to pre-event levels ahead of the CIE (Fig. 6-2). This suggests that nannoplankton can adapt to changing environments and then resume normal productivity once they adjust to the environmental stresses, although the possibility of a lag in the ocean’s response caused by system inertia cannot be excluded.

Summary

New and published data confirm surface-water oligotrophy in the southern high-latitude open ocean across the PETM. Depletion of nutrients culminated during the peak of the event, which may have resulted from intensified water-column stratification. Rapid ocean acidification as a result of elevated pCO2 from the onset to the peak of the event caused a major drop in the accumulation of dissolution-susceptible holococcoliths. Similarities and contrasts between the results of this study and previous work from open- ocean sites (ODP 690 and 1209) and shelf margins (e.g., Site 1259 and Wilson Lake, New Jersey) further demonstrate that the response to the PETM was consistent in the open oceans, but could be localized on the continental shelves where nutrient regimes depend on the local geologic setting and oceanographic conditions.

Acknowledgements

We thank James Arney for his preliminary delineation of the PETM interval we examined at Site 1135, and Yang Wang and Yingfeng Xu for facilitating our analyses of bulk foraminiferal isotopic measurements. Alan Cooper, Timothy Bralower, and John Barron provided constructive and helpful comments that greatly improved this manuscript. This research used samples and

133 data provided by ODP. Funding provided by JOI USSAC (for ODP Legs 183 and 207) and NSF- OPP grants 0126218 and 0230469 to SWW.

134

Figure 6-1. ODP Site 1135 (★ ) and other reference sites (●) on an ODSN 55-Ma plate reconstruction.

135

Figure 6-2. Percentage abundances of nannofossil groupings and bulk foraminiferal δ13C and δ18O across the PETM from ODP Site 1135. K-selected taxa include Discoaster, Fasciculithus, and Sphenolithus, while r-selected taxa include Biscutum, Chiasmolithus, and Hornibrookina. Age assignments followed Zachos et al. (2003) based on the apparent duration of the CIE. Bulk density data are from the ODP Janus Database. Variations in percentage of unidentifiable isolated nannofossil shields generally depict the preservation of fossil assemblages. Note different abundance scales. mbsf = meters below sea floor.

136 Table 6-1. Calcareous nannofossil counts in ODP Hole 1135A.

137

APPENDIX A

Calcareous nannofossils considered in this paper in alphabetical order of generic epithets.

Amaurolithus amplificus (Bukry & Percival, 1971) Gartner & Bukry, 1975 Amaurolithus primus (Bukry & Percival, 1971) Gartner & Bukry, 1975 Amaurolithus ninae Perch-Nielsen, 1977 Amaurolithus tricorniculatus (Gartner, 1967) Gartner & Bukry, 1975 Calcidiscus macintyrei (Bukry & Bramlette, 1969) Loeblich & Tappan, 1978 Calcidiscus premacintyrei Theodoridis, 1984 Catinaster coalitus Martini & Bramlette, 1963 Catinaster mexicanus Bukry, 1971 Ceratolithus acutus Gartner & Bukry, 1974 Ceratolithus armatus Müller, 1974 Ceratolithus cristatus Kamptner, 1950, emend. Bukry & Bramlette 1968 Ceratolithus rugosus Bukry & Bramlette, 1968 Ceratolithus telesmus Norris, 1965 Coccolithus miopelagicus Bukry, 1971, emend. Wise, 1973 Coccolithus pelagicus (Wallich, 1877) Schiller, 1930 Coronocyclus nitescens (Kamptner, 1963) Bramlette & Wilcoxon, 1967 Cyclicargolithus abisectus (Müller, 1970) Wise, 1973 Cyclicargolithus floridanus (Hay et al., 1967) Bukry, 1971 Discoaster adamanteus Bramlette & Wilcoxon, 1967 Discoaster asymmetricus Gartner, 1969 Discoaster aulakos Gartner, 1967 Discoaster bellus Bukry & Percival, 1971 Discoaster berggrenii Bukry, 1971 Discoaster blackstockae Bukry, 1973

138 Discoaster bollii Martini & Bramlette, 1963 Discoaster braarudii Bukry, 1971 Discoaster brouweri TAN SIN HOK, 1927, emend. Bramlette & Riedel, 1954 Discoaster calcaris Gartner, 1967 Discoaster challengeri Bramlette & Riedel, 1954 Discoaster decorus (Bukry, 1971) Bukry, 1973 Discoaster deflandrei Bramlette & Riedel, 1954 Discoaster exilis Martini & Bramlette, 1963 Discoaster hamatus Martini & Bramlette, 1963 Discoaster icarus Stradner, 1973 Discoaster intercalaris Bukry, 1971 Discoaster kugleri Martini & Bramlette, 1963 Discoaster loeblichii Bukry, 1971 Discoaster mendomobensis Wise, 1973 Discoaster micros (Theodoridis, 1984) Kaenel & Villa, 1996 Discoaster moorei Bukry, 1971 Discoaster neohamatus Bukry & Bramlette, 1969 Discoaster neorectus Bukry, 1971 Discoaster pansus (Bukry & Percival, 1971) Bukry, 1973 Discoaster pentaradiatus TAN SIN HOK, 1927, emend. Bramlette & Riedel, 1954 Discoaster perclarus Hay in Hay et al., 1967 Discoaster prepentaradiatus Bukry & Percival, 1971 Discoaster pseudovariabilis Martini & Worsley, 1971 Discoaster quinqueramus Gartner, 1969 Discoaster sanmiguelensis Bukry, 1981 Discoaster surculus Martini & Bramlette, 1963 Discoaster tristellifer Bukry, 1976 Discoaster variabilis Martini & Bramlette, 1963 Emiliania huxleyi (Lohmann, 1902) Hay & Mohler in Hay et al., 1967 Gephyrocapsa caribbeanica Boudreaux & Hay, 1967 Gephyrocapsa oceanica Kamptner, 1943

139 Gephyrocapsa Kamptner, 1943 (small) Hayaster perplexus (Bramlette & Riedel, 1954) Bukry, 1973 Helicosphaera burkei Black, 1971 Helicosphaera carteri (Wallich, 1877) Kamptner, 1954 Helicosphaera euphratis Haq, 1966 Helicosphaera ٛ artiniٛٛ (Bukry & Percival, 1971) Jafar & Martini, 1975 Helicosphaera ٛ artini Gartner, 1980 Helicosphaera kamptneri (Hay & Mohler in Hay et al., 1967) Locker, 1973 Helicosphaera sellii (Bukry & Bramlette, 1969) Jafar & Martini, 1975 Helicosphaera walbersdorfensis Mueller, 1974 Helicosphaera wallichii (Lohmann, 1902) Boudreaux & Hay, 1969 Pontosphaera japonica (Takayama, 1967) Nishida, 1971 Pontosphaera multipora (Kamptner, 1948) Roth, 1970, emend. Burns, 1973 Pontosphaera versa (Bramlette & Sullivan, 1961) Sherwood, 1974 Pseudoemiliania ٛ artiniٛ (Kamptner, 1963) Gartner, 1969 Reticulofenestra gelida (Geitzenauer, 1972) Backman, 1978 Reticulofenestra haqii Backman, 1978 Reticulofenestra [= Dictyococcites] minuta Roth, 1970 Reticulofenestra minutula (Gartner, 1967) Haq & Berggren, 1978 Reticulofenestra perplexa (Burns, 1975) Wise, 1983 Reticulofenestra producta (Kapmtner, 1963) Wei & Thierstein, 1991 Reticulofenestra pseudoumbilica (Gartner, 1967) Gartner, 1969 Reticulofenestra rotaria Theodoridis, 1984 Rhabdosphaera clavigera Murray & Blackman, 1898 Rhabdosphaera procera Martini, 1969, emend. Jafar, 1975 Scapholithus fossilis Deflandre in Deflandre & Fert, 1954 Scyphosphaera Lohmann, 1902 Sphenolithus abies Deflandre in Deflandre and Fert, 1954 Sphenolithus compactus Backman, 1980 Sphenolithus heteromorphus Deflandre, 1953 Sphenolithus moriformis (Brönnimann and Stradner, 1960) Bramlette and Wilcoxon, 1967

140 Sphenolithus neoabies Bukry and Bramlette, 1969 Sphenolithus verensis Backman, 1978 Syracosphaera pulchra Lohmann, 1902 Triquetrorhabdulus rugosus Bramlette & Wilcoxon, 1967 Umbellosphaera irregularis Paasche in Markali and Paasche, 1955 Umbilicosphaera jafari Müller, 1974 Umbilicosphaera rotula (Kamptner, 1956) Varol, 1982 Umbilicosphaera sibogae (Weber-van Bosse, 1901) Gaarder, 1971

141

APPENDIX B

Calcareous nannofossils considered in this paper in alphabetical order of generic epithets. Biscutum sp. Plate 4-2, Figs. 33-36. Braarudosphaera bigelowii (Gran & Braarud, 1935) Deflandre, 1947 Campylosphaera dela Hay & Mohler, 1967. Plate 4-2, Figs. 21-25. Campylosphaera eodela Bukry & Percival, 1971. Plate 4-2, Figs. 26-27. Chiasmolithus bidens (Bramlette & Sullivan, 1961) Hay & Mohler, 1967. Plate 4-2, Figs. 1-2. Chiasmolithus (Sullivania) consuetus (Bramlette & Sullivan) Hay & Mohler, 1967. Plate 4-2, Figs. 3-5. Chiasmolithus solitus (Bramlette & Sullivan, 1961) Locker, 1968. Plate 4-2, Fig. 6. Chiasmolithus [= Sullivania] californicus (Sullivan, 1964) Hay & Mohler, 1967. Plate 4-2, Figs. 7-9. Chiasmolithus sp. Coccolithus formosus (Kamptner, 1963) Wise, 1973 Coccolithus pelagicus (Wallich, 1877) Schiller, 1930. Plate 4-1, Figs. 35-36. Coccolithus subpertusus (Hay & Mohler, 1967) Wei & Pospichal, 1991 Coronocyclus prionion (Deflandre & Fert, 1954) Stradner & Edwards, 1968. Plate 4-2, Fig. 40. Cruciplacolithus cribellum (Bramlette & Sullivan, 1961) Romein, 1979 Cruciplacolithus latipons Romein, 1979. Plate 4-2, Fig. 11. Cruciplacolithus sp. Plate 4-2, Fig. 10. Discoaster anartios Bybell & Self-trail, 1995 Discoaster araneus Bukry, 1971. Plate 4-1, Figs. 1-5, 11-21. Discoaster diastypus Bramlette & Sullivan, 1961 Discoaster falcatus Bramlette & Sullivan, 1961 Discoaster cf. D. falcatus Bramlette & Sullivan, 1961 Discoaster lenticularis Bramlette & Sullivan, 1961. Plate 4-2, Fig. 16.

142 Discoaster mohleri Bukry & Percival, 1971 Discoaster cf. D. mohleri Bukry & Percival, 1971 Discoaster multiradiatus Bramlette & Riedel, 1954. Plate 4-2, Figs. 13-15. Discoaster salisburgensis Stradner, 1961 Discoaster sp. Plate 4-2, Fig. 17. Ellipsollithus distichus (Bramlette & Sullivan, 1961) Sullivan, 1964. Plate 4-2, Fig. 38. Ellipsolithus macellus (Bramlette & Sullivan, 1961) Sullivan, 1964. Plate 4-2, Fig. 37. Fasciculithus involutus Bramlette & Sullivan, 1961 Fasciculithus sidereus Bybell & Self-Trail, 1995. Plate 4-2, Fig. 20. Fasciculithus sp. Fasciculithus thomasii Perch-Nielsen, 1971. Plate 4-2, Fig. 19. Fasciculithus tympaniformis Hay & Mohler, 1967. Plate 4-2, Fig. 18. Hornibrookina arca Bybell & Self-trail, 1995. Plate 4-2, Figs. 31-32. Lophodolithus nascens Bramlette & Sullivan, 1961 Markalius apertus Perch-Nielsen, 1979. Plate 4-1, Figs. 39-40. Markalius inversus (Deflandre in Deflandre & Fert, 1954) Bramlette & Martini, 1964. Plate 4-1, Figs. 37-38. Markalius cf. M. apertus Perch-Nielsen, 1979. Plate 4-1, Figs. 6-9, 22-26. Neochiastozygus concinnus (Martini, 1961) Perch-Nielsen, 1971 Neochiastozygus distentus (Bramlette & Sullivan, 1961) Perch-Nielsen, 1971 Neochiastozygus junctus (Bramlette & Sullivan, 1961) Perch-Nielsen, 1971. Plate 4-2, Fig. 28. Neococcolithes dubius (Deflandre in Deflandre & Fert, 1954) Black, 1967 Neococcolithes protenus (Bramlette & Sullivan, 1961) Black, 1967. Plate 4-2, Fig. 29. Placozygus sigmoides (Bramlette & Sullivan, 1961) Romein, 1979. Plate 4-2, Fig. 30. Pontosphaera scissura (Perch-Nielsen, 1971) Romein, 1979. Plate 4-2, Fig. 42. Pontosphaera sp. Rhabdosphaera tenuis Bramlette & Sullivan, 1961 Rhomboaster cuspis Bramlette & Sullivan, 1961 Rhomboaster spineus (Shafik & Stradner, 1971) Perch-Nielsen, 1984. Plate 4-1, Fig. 28. Rhomboaster sp. Plate 4-1, Fig. 27. Scapholithus fossilis Deflandre in Deflandre & Fert, 1954

143 Scapholithus rhombiformis Hay & Mohler, 1967. Plate 4-2, Fig. 41. Sphenolithus anarrhopus Bukry and Bramlette, 1969 Sphenolithus moriformis (Brönnimann and Stradner, 1960) Bramlette & Wilcoxon, 1967 Thoracosphaera sp. Plate 4-2, Fig. 39. Toweius eminens (Bramlette & Sullivan, 1961) Perch-Nielsen, 1971. Plate 4-1, Figs. 31-32. Toweius cf. eminens. Plate 4-1, Fig. 33. Towieus occultatus (Locker, 1967) Perch Nielsen, 1971 Toweius pertusus (Sullivan, 1965) Romein, 1979. Plate 4-1, Fig. 34. Toweius cf. T. selandianus Perch-Nielsen 1979. Plate 4-1, Fig. 30. Toweius sp. Plate 4-1, Fig. 29. Transversopontis sp. Tribrachiatus bramlettei (Brönnimann & Stradner, 1960) Proto Decima et al., 1975. Plate 4-1, Fig. 10. Zygodiscus adamas Bramlette & Sullivan, 1961 Zygrhablithus bijugatus (Deflandre in Deflandre & Fert, 1954) Deflandre, 1959

144

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BIOGRAPHICAL SKETCH

EDUCATION China University of Geosciences Geology M.S., 1996 China University of Geosciences Geology B.S., 1993

PROFESSIONAL EXPERIENCE: 2002-Present Teaching and research assistant at the Florida State University 2000-2001 Manager of Research Administration, Nanhai East Institute (NHEI), China National Offshore Oil Corporation (CNOOC) 1998-2000 Manager of Paleontology Group at Science & Technology Research Center, China Offshore Oil Nanhai East Corp., CNOOC 1996-1998 Micropaleontologist at Science & Technology Research Center, China Offshore Oil Nanhai East Corp., CNOOC Research Cruise: 2002.11-2003.1 Shipboard scientist on JOIDES Resolution, ODP Leg 206 Research Participation: 1996-2001 Neogene biostratigraphy and sequence stratigraphy of Pearl River Mouth Basin, northern shelf of the South China Sea 1994-1996 Paleoceanographic study of the Bashi Strait and the South China Sea during the past 20,000 years 1994-1995 Geological engineering study on the dam site of Shuibuya hydrodynamic power station, Qingjiang, west Hubei Province 1992-1993 Tertiary paleooceanographical study of Qinling, China

Awards: • 10th International Symposium on Antarctic Earth Science (ISAES) travel grant: 2007, $1200

170 • Florida State University Conference Presentation Grant: 2007, $300; 2006 (twice), $600 • U.S. Scientific Advisory Committee (USSAC) grant (Task Order F001790): 2003, $40157 • Third-class honor in Scientific Progress of China Offshore Oil Nanhai East Corp.: 1998 • Model Graduate Cadre at university level of China University of Geosciences: 1994, 1995 • Model Student of China University of Geosciences: 1992 • Model Student of Geology Department, China University of Geosciences: 1991

TEACHING EXPERIENCE: • Dynamic Earth Lab (GLY1000L), Physical Geology Lab (GLY2010C), and Paleontology lab (GLY3610C).

RESEARCH INTERESTS: • Taxonomy, biostratigraphy, and evolution of microfossils (e.g., calcareous nannofossils, foraminifers, and spores and pollen), and their applications to Cenozoic earth history, sequence stratigraphy, and oil and gas exploration; • Stable isotope geochemistry and applications in paleoenvironmental reconstruction • Paleoceanography, paleogeography, and paleoclimatology by means of microfossils and geochemical proxies (e.g., stable C and O isotopes, Mg/Ca ratios, Uk’37); • Multivariate statistical analyses of geological data.

SYNERGISTIC ACTIVITIES: Membership of AGU since 2004 Membership of International Nannoplankton Association (INA) since 2003

PUBLICATIONS: Peer-reviewed papers: • Jiang, S., and Wise, S.W. Taxonomic note: a new Coccolithus species thriving during the Paleocene/Eocene Thermal Maximum. Submmited to Journal of Nannoplankton Research. • Chen, Z., Jiang, S., Huang, C-Y, Yan, W., Lu, J., and Chen, M-H. Concentrations and source of heavy metals in surface sediments of Nansha Islands and adjacent sea area in the South China Sea. Submitted to Continental Shelf Research.

171 • Jiang, S., and Wise, S.W., 2007. Abrupt turnover in calcareous-nannoplankton assemblages across the Paleocene/Eocene Thermal Maximum: implications for surface- water oligotrophy over the Kerguelen Plateau, Southern Indian Ocean. In Antarctica: A Keystone in a Changing World – Online Proceedings of the 10th ISAES, edited by A. K. Cooper and C. R. Raymond et al., USGS Open-File Report 2007-1047, Short Research Paper 024, 5 p.; doi:10.3133/of2007-1047.srp024. • Jiang, S., Wise, S.W., Jr., and Wang, Y., 2007. Cause of the middle/late Miocene carbonate crash: dissolution or low productivity? In Teagle, D.A.H., Wilson, D.S., Acton, G.D., and Vanko, D.A. (Eds.), Proc. ODP, Sci. Results, 206: College Station, TX (Ocean Drilling Program), 1–24. doi:10.2973/odp.proc.sr.206.013.2007. • Jiang, S., and Wise, S.W., Jr., 2007. Upper Cenozoic calcareous nannofossil biostratigraphy and inferred sedimentation, ODP Leg 206, East Pacific Rise. In Teagle, D.A.H., Wilson, D.A., Acton, G.A., and Vanko, D.A. (Eds.), Proc. ODP, Sci. Results, 206: College Station, TX (Ocean Drilling Program), 1–25. doi:10.2973/odp.proc.sr.206.008.2007 • Jiang S., and Wise, S.W., 2006. Surface-water chemistry and fertility variations in the tropical Atlantic across the Paleocene/Eocene Thermal Maximum as evidenced by calcareous nannoplankton from ODP Leg 207, Hole 1259B. Revue de micropaléontologie, 49(4): 227–244. • Jiang, S., and Wise, S.W., 2006. Taxonomic note: Discoaster stellulus Gartner, 1967, emended. Journal of Nannoplankton Research, 28(2): 81–83. • Lin, J., Zhang, J., Jiang, S., Wang, S., Xu, B., and Wei, M., 2004. The Neogene foraminiferal stratigraphy of the LH19-4-1 bore hole, Pearl Rive Mouth Basin, South China Sea. Journal of Stratigraphy (Chinese), 28(2): 120–125. • Persico, D., Wise, S.W. Jr., and Jiang, S., 2003. Oligocene-Holocene calcareous nannofossil biostratigraphy and diagenetic etch patterns on Quaternary placoliths at ODP Site 1139 on Skiff Bank, Northern Kerguelen Plateau. In Frey, F.A., Coffin, M.F., Wallace, P.J., and Quilty, P.G. (Eds.), Proc. ODP, Sci. Results, 183 [Online]. Available from World Wide Web: .

172 • Jiang S., 1998. High-resolution calcareous nannofossil biostratigraphy and sequence stratigraphy of Well PY27-2-1, Pearl River Mouth Basin, China. China Offshore Oil and Gas (Geology), 1999, 13(3): 189-195. • Duan W., Chen F., Wei K., and Jiang, S., 1998. Calcareous nannofossils in the northeastern South China Sea since the last 15 ka and their significance to paleoceanography: Haiyang Dizhi yu Disiji Dizhi = Marine Geology & Quaternary Geology, 18(1), p. 13-22, (Chinese) (English sum.), illus., 18 ref., February 1998.

Abstracts and Presentations: • Jiang, S., and Wise, S.W., 2006. Surface-water chemistry and fertility variations in the tropical Atlantic across the Paleocene/Eocene thermal maximum as evidenced by calcareous nannoflora from Demerara Rise, central Atlantic, International Nannoplankton Association 11th Conference, September 2006, Lincoln, NE, USA, Abstract and oral presentation. • Jiang, S., Wise, S.W., and Wang, Y., 2006. Persistent El Niño-like conditions during the middle/late Miocene carbonate crash. EOS Trans. AGU, 87(36), West. Pac. Geophys. Meet. Suppl., Abstract V44B-03 plus oral presentation. • Foley, S., Ingram, W., Jiang, S., Stant, A., and Wise, S.W., 2004. Oligocene-middle Miocene calcareous nannofossils from the northernmost Kerguelen Plateau (46 degrees S); warm-water indicators and influences. Abstract, Journal of Nannoplankton Research, vol.26, no.2, pp.43.

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